STRUCTURAL GEOLOGY AND TECTONICS OF THE PALEOPROTEROZOIC ROCKS OF THE MOUNT RUSHMORE QUADRANGLE, BLACK HILLS, SOUTH DAKOTA ______
A Dissertation
presented to the Faculty of the Graduate School University of Missouri-Columbia
______
In Partial Fulfillment Of the Requirements for the Degree
Doctor of Philosophy
______
By
JOSEPH CHRISTOPHER HILL
AUGUST 2006
Dr. Peter I. Nabelek, Advisor Dr. Robert L. Bauer, Co-Advisor
To my son,
Aidan Christopher
You gave me the reason to persevere…
And my wife,
Kristin Leigh
You gave me the strength.
Nothing is so firmly believed as what we least know.
- Michel de Montaigne
ACKNOWLEDGEMENTS
It has been a long, bumpy ride over a long, long road and yet somehow I have arrived at the end of it – definitely a different person from when I started.
As Henri-Frédéric Amiel said: “All appears to change when we change.” I have
learned a lot, been given the opportunity to receive an excellent education, and in
the process learned a lot about myself. I didn’t enjoy every minute of it, I regret
some of it, and I understand a little better now that it is all over why kismet chose
this path. To paraphrase Thomas Jefferson, the harder I work the luckier I get.
I have lots of people to thank for guiding me, pushing me, and suffering
me. First I would like to thank the members of my committee: Drs. Glen
Himmelberg (emeritus), Erik Loehr, Kevin Shelton, and Alan Whittington and
especially my co-advisors Drs. Robert Bauer and Peter Nabelek. I’m sure we
would all agree that we wouldn’t want to do it all over again. A big thanks to
Linda Garrison and Marsha Huckabey for everything they do to keep the
Geological Sciences department going. Thank you to the Dr. Tom Freeman for
being such a positive influence on me and for offering such good advice.
I would also like to thank all the great friends that I’ve had while at Mizzou
and point out that Aaron Johnson, Damon Bassett, Tim Huff, and Cathy Zumsteg
are the salt of the earth, and that Aaron and I are not allowed out with each other
without supervision as our wives won’t post bail. I would like to thank Brian
Fagnan for just being the “Dude” and for hanging out all those summers in the
Black Hills, talking geology, drinking copious amounts of alcoholic beverages,
and laughing a lot. I will see you at the 75th Sturgis rally. Drew and Rebecca
ii Hilpert became our closest friends while here in Missouri and for no other reason than that, I think our time in Missouri was well spent. As I write this, I find I should attach another bibliography to this work of names of friends too numerous to mention, I hope you know who you are (Carolina, Chris, John, Angie, Eric, and the rest). I wish each and every one of you the very best.
Finally, to my wife Kristin, you know I owe this to your love and patience.
To us babe, and to the fulfillment of a promise.
iii TABLE OF CONTENTS
ACKNOWLEDGEMENTS ...... ii
LIST OF ILLUSTRATIONS...... vi
LIST OF TABLES ...... x
LIST OF PLATES ...... xi
ABSTRACT ...... xii
INTRODUCTION...... 1
Importance of the Mount Rushmore Quadrangle as a Proxy for Understanding the Tectonics of the Black Hills ...... 5
Research Objectives...... 9
Regional Geologic Setting...... 10
Archean Basement...... 11
Paleoproterozoic Rocks...... 13
~2550 – 2480 Ma Rift Succession...... 13
~2015 – 1885 Ma Continental Slope-Rise Type Succession ...... 15
1715 Ma Harney Peak Granite ...... 15
Deformation and Metamorphism ...... 16
Tectonic Models ...... 18
Unresolved Issues of Black Hills Geology...... 20
Lithotectonic Units versus Stratigraphy ...... 21
Problems with Stratigraphic Correlation ...... 24
Previous Stratigraphic Correlations ...... 26
Methodology...... 27
Previous Geologic Mapping ...... 28 STRUCTURAL RELATIONSHIPS AND DEFORMATIONAL HISTORY...... 29
Summary of Previous Structural Studies...... 30
Deformational Events and Features Considered in this Study...... 37
Bedding, Primary Sedimentary Structures, and Compositional Variation...... 38
D1 Deformation ...... 42
Foliation ...... 43
Folds...... 44
D2 Deformation ...... 47
Foliation ...... 48
Variations in S2 and Related Structures Across the Study Area...... 53
Folds...... 57
Lineations ...... 59
Variation of D2 Deformation Features by Fault Block ...... 63
Faults...... 65
Recognition of a Major Structural Discontinuity...... 71
D3 Deformation ...... 75
D4 Deformation ...... 76
Foliation ...... 77
Faulting...... 79
D5 Deformation ...... 80
Foliation ...... 80
Folds...... 81
Summary of Deformational Events...... 81
TECTONIC OVERVIEW AND OBLIQUE CONVERGENCE MODEL ...... 87 v Laurentian Assembly: Overview...... 89
Incongruities with Laurentian Assembly ...... 90
Problems with”Trans-Hudson”...... 91
Exposure...... 92
Geophysical Interpretation...... 93
Age and Character of “Trans-Hudson” Rocks ...... 97
Wyoming Province Orogenies ...... 100
Transpressional Model for the Study Area ...... 101
Structural Discontinuity...... 102
Emplacement of the Harney Peak Granite ...... 104
40 39 M1 Porphyroblasts and Ar/ Ar Data...... 110
Overall Structural Pattern and “Cross Folds”...... 112
Application of the Transpressional Model to the Black Hills ...... 114
Transpressional Model...... 115
Extrapolation to the Tectonics of the Wyoming Province ...... 117
CONCLUSIONS ...... 120
REFERENCES CITED...... 125
APPENDIX A: Lithostratigraphic Units ...... 139
APPENDIX B: Lithotectonic Units Versus Stratigraphy...... 143
APPENDIX C...... 160
VITA ...... 161
vi LIST OF ILLUSTRATIONS
Figure 1 – Precambrian tectonic elements of Laurentia...... 2
Figure 2 – Tectonic setting and location of the Black Hills uplift...... 3
Figure 3 – Location and generalized geologic map of the Precambrian Black Hills...... 7
Figure 4 – Geologic map of the Mount Rushmore 7.5-minute quadrangle...... 8
Figure 5 – Outline map of fault bounded blocks delineated during this study in the Mt. Rushmore 7.5-minute quadrangle...... 23
Figure 6 – Metamorphic zones of the Precambrian Black Hills ...... 32
Figure 7 – Generalized geologic map of the Precambrian Black ...... 33
Figure 8 – Typical exposure of metasedimentary rocks...... 39
Figure 9 – Atypical exposures of the typical relationship of bedding (S0) to compositional banding (Sc) in metapsammites within the study area ...... 40
Figure 10 – Example of bedding recognition through petrographic analysis...... 41
Figure 11 – Photograph of early recumbent fold of interbedded metagraywacke and micaceous schist...... 46
Figure 12 – Photographs of F2 fold in mica-quartz schists showing S2 axial planar cleavage...... 49
Figure 13 – Outcrop photographs of F2 fold showing S2 axial planar cleavage in folded mica-quartz schists...... 50
Figure 14 – Photomicrographs of S2 foliation demonstrating the variable nature of S2...... 51
Figure 15 – Lower hemisphere equal area projections of contoured poles to S2 foliation and axial-planes of F2 Folds – Fault-blocks A - C...... 54
Figure 16 – Lower hemisphere equal area projections of contoured poles to S2 foliation and axial-planes of F2 Folds – Fault-blocks D - G ...... 55
Figure 17 – Lower hemisphere equal area projections of contoured poles to S2 foliation and axial-planes of F2 Folds – Fault-blocks H - I ...... 56
vii Figure 18 – Lower hemisphere equal area projections of poles to F2 fold axial planes measured during this study ...... 58
Figure 19 – Lower hemisphere equal area projections of L2 lineations and F2 hinge- lines measured during this study...... 60
Figure 20 – Lower hemisphere equal area projections of L2 lineations (contoured) and F2 hinge-lines (black squares)...... 61
Figure 21 – Excellent example of L2 lineations defined by elongate quartz minerals in quartz-mica schist ...... 62
Figure 22 – Breccia in quartzite formed during early(?) brittle faulting in the study area ...... 66
Figure 23 – Expression of the Keystone Fault in metagraywacke...... 67
Figure 24 – Photomicrograph of graphitic rocks from station MR270 showing left- lateral shear as defined by the asymmetry of the porphyroclasts ...... 68
Figure 25 – Photomicrograph of graphitic metasiltstone from station MR349 demonstrating left-lateral shear as defined by the asymmetry of the mica-quartz aggregates...... 69
Figure 26 – Photomicrograph of sample MR675, a sheared metapelite collected near the Silver City fault in fault-block H ...... 70
Figure 27 – Outline map of the Precambrian Black Hills uplift showing locations of major faults mentioned in this study...... 73
Figure 28 – Lower hemisphere equal area projections of poles to S0/Sc foliation in fault-block B...... 78
Figure 29 – Lower hemisphere equal area projection of poles to S4 foliation measured in fault-block A...... 79
Figure 30 – Field photographs showing the well developed S5 spaced cleavage and its relationship to the dominant S2 fabric ...... 82
Figure 31 – Field photograph of S5 crenulation cleavage developed east of the Silver City fault ...... 83
Figure 32 – Photomicrographs of S5 as defined by bounding surfaces of microlithons ...... 84
Figure 33 – Photomicrographs of S5 foliation ...... 85
Figure 34 – Photomicrograph of sample MR915 showing well developed S5 crenulation cleavage trending subvertically right to left...... 86 viii Figure 35 – Photograph of possible F5 fold showing warping of S2 foliation ...... 86
Figure 36 – Tectonic crustal model of Klasner and King (1990)...... 96
Figure 37 – COCORP seismic transect across the southern Trans-Hudson orogen beneath the Williston Basin...... 96
Figure 38 – Location of the NACP in Central North America ...... 97
Figure 39 – Generalized geologic map of the northeastern Black Hills with scaled inset of detailed mapping of Mt. Rushmore quadrangle ...... 104
Figure 40 – Diagram showing map view of mechanics of a releasing bend step-over along a sinistral strike-slip fault ...... 108
Figure 41 – Block diagram of a negative flower structure ...... 109
Figure 42 – Possible configuration of the Wyoming Province and Dakota block during 1760-1715 Ma collision ...... 117
Figure 43 – Depiction of potential configuration of major cratonic blocks of now part of southern Laurentia ...... 119
ix
LIST OF TABLES
Table 1 – Summary of lithostratigraphic units used in this study ...... 24
Table 2 – Summary of Proterozoic deformational events and related structures ..... 36
Table 3 – Revised summary of deformational events and related structures based on this study ...... 88
Table 4 – Summary of 40Ar/39Ar results from the northern Black Hills, Dahl et al. (1999a)...... 112
x
LIST OF PLATES
Plate I – Geologic Map of the Mount Rushmore 7.5-minute Quadrangle
Plate II – Station Locations
xi ABSTRACT
The Paleoproterozoic assembly of Laurentia, the cratonic core of North
America, involved the assembly of numerous Archean microcontinents, between
~ 1.96 to 1.70 Ga, and continued with the accretion of island arc terranes until about 1.63 Ga. The Black Hills, South Dakota lie along the eastern margin of the
Wyoming Archean craton, one of the assembled Archean microcontinents, and
the western edge of the southern projection of the Paleoproterozoic Trans-
Hudson orogen (THO), which formed as a result of the collision of the Superior
Archean craton and Archean cratons to the west. Most of the southern THO is
covered by a thick sequence of Phanerozoic sedimentary rocks; however, the
Black Hills contain an exposed core of Precambrian rocks that are critical to our understanding of large-scale and complex aspects of the assembly of southern
Laurentia.
The Precambrian rocks of the Black Hills uplift record multiple deformational and metamorphic events related to three orogenic events that have affected the study area: (1) the ~1780 Ma Black Hills orogeny, (2) the ~ 1720 Ma
Dakotan orogeny, and (3) the ~1690 Ma Central Plains orogeny. In response to these three orogenies, the rocks of the Mount Rushmore quadrangle, which are the focus of this study, have undergone a complex Paleoproterozoic history of ductile folding, intrusion of the Harney Peak granite (HPG)(ca. 1715 Ma), and two period of metamorphism.
xii The study area was divided into nine fault-bounded blocks that have
different structural histories and were, therefore, evaluated as distinct structural domains. A major structural discontinuity defined by the Keystone and Empire
Mine faults separates six of the structural domains, to the east of the discontinuity, from three structural domains to the west. Evidence from detailed mapping demonstrates that this discontinuity is a sinistral strike-slip fault zone
that truncates all earlier structures to the southwest. The domains to the west of
the discontinuity are proximal to the NE margin of the HPG but differ considerably
in their response to the emplacement of the HPG. The domains to the east of the
discontinuity are interpreted as being part of an allochthonous group of fault
blocks with an overall sense of sinistral displacement. The allochthonous blocks
northeast of the mapped discontinuity are believed to represent a series of
“tectonic slivers” that moved unknown distances along the eastern margin of the
Archean Wyoming craton.
Structural features mapped in the study area and the overall structural
trends of the east-central Black Hills may be best explained by an oblique
convergence model involving oblique collision of an intervening continental
fragment, the Dakota block, during terminal Wyoming-Superior collision. This
collision resulted in early east-west convergence that produced the earliest folds,
faults, and metamorphism recognized in the Black Hills. During the later stages of
collision, the overall convergence became more northwest directed, resulting in
the strong overprint of NW-trending folds and fabrics. The series of anastomosing
strike-slip faults evident in the east-central Black Hills were formed as a result of
xiii sinistral transpression when the Dakota block slipped along the eastern margin of
Wyoming craton.
Emplacement of the Harney Peak Granite and the associated second
prograde metamorphic event occurred during the waning stages deformation as dilatational space was created in a releasing bend during north directed transpression. Uplift of the fault-bounded block containing the Harney Peak
Granite is likely a result of differential buoyancy related to the emplacement of the melt coupled with late compression of the releasing bend reactivating the faults in high-angle reverse motion. Movement along strike-slip faults continued after the emplacement of the Harney Peak Granite. A late deformational event affects the southern half of the study area and is likely a result of far-field stresses related to the Central Plains orogen ~ 1690 Ma.
xiv INTRODUCTION
The assembly of Laurentia, the cratonic core of North America, was
initiated during Early Proterozoic time by the assembly of numerous Archean
microcontinents, between ~ 1.96 to 1.81 Ga, and continued with the accretion of
island arc terranes until about 1.63 Ga (Hoffman, 1988). As a result of this
process, Laurentia is composed of a network of Early Proterozoic orogenic belts,
accreted island arc terranes, and formerly independent Archean microcontinents with highly deformed margins (Hoffman, 1988) (Figure 1). These elements have remained coherent and relatively rigid for at least the last one billion years.
However, the nature and timing of many of the Paleoproterozoic events involved in the assembly of Laurentia remain poorly understood. The keys to a better understanding of this process lie in unraveling the complex deformation and thermal events produced along the collisional boundaries of the various assembled components of Laurentia.
The Black Hills of South Dakota lie along the eastern margin of the
Archean Wyoming craton, one of the assembled Archean microcontinents, and the western edge of the southern projection of the Early Proterozoic Trans-
Hudson orogen (Figure 2). They contain the only significant exposures of
Precambrian rocks in a region where most Precambrian units are covered by a thick sequence of Phanerozoic sedimentary rocks. As a result, the Black Hills contain Precambrian rock exposures that are critical to our understanding of large-scale and complex aspects of the assembly of southern Laurentia, and they
Figure 1 - Precambrian tectonic elements of Laurentia relative to modern North America. Archean microcontinents include the Slave craton, Hearne-Rae province, Wyoming craton, Superior craton, Medicine Hat block (MHB). Darker blue represents Paleoproterozoic orogenic events. See text for more discussion. Modified from Hoffman (1988). Not labeled on map are the Thelon Orogen (Slave/Hearne-Rae collision), Vulcan Tectonic zone (MHB northern boundary), Great Falls Tectonic Zone (MHB southern boundary), and the Torngat Orogen (Nain/Superior collision).
2
Figure 2 - Tectonic setting and location of the Black Hills uplift showing the major tectonic elements of the mid-continent region of North America. Note the Laramide basement cored uplifts (dark blue) that define the Wyoming Province. The Cheyenne Belt/Green Mountain block is shown in green. The southern Yavapai province marks the boundary of 1.7 Ga basement. Modified from Karlstrom et al. (2002).
3 offer a unique opportunity to study part of the 1900 Ma to 1600 Ma
Paleoproterozoic orogenic events involved in this assembly. Specifically, the
Black Hills offer the only surface exposure, south of 54o N latitude, of the deformed belt of Archean and Paleoproterozoic rocks in the collisional zone between the Archean Wyoming and Superior provinces.
Based largely on seismic reflection and refraction (e.g. Baird et al., 1995,
1996; Lucas et al., 1993, 1994; Nelson et al., 1993; White et al., 1994; Latham et al., 1988), magnetotelluric (e.g., Jones et al., 1993, 1995; Garcia and Jones,
2005), electromagnetic (e.g. Jones et al., 2005 and references therein; Toft et al.,
1993), geomagnetic depth sounding (e.g., Jones et al., 2005; Camfield and
Gough, 1977; Camfield et al., 1970), paleomagnetic (e.g., Stauffer, 1984) and gravity (e.g. Klasner and King, 1986, 1990) geophysical evidence, this collisional zone has been broadly interpreted to be a southern extension of the Trans-
Hudson orogeny that was truncated by a collage of island-arc terranes of the ca.
1680 Ma Central Plains orogen (Klasner and King, 1990) (Figure 2). However, differences between the Trans-Hudson events recorded in Canada (1850 - 1800
Ma; Bickford et al., 1990) and events recorded in the Black Hills (1775 – 1690
Ma; Dahl et al., 2005a,b) allows for multiple hypotheses regarding the assemblage of southern Laurentia. The Canadian Trans-Hudson is a collage of mainly juvenile, arc-related Paleoproterozoic domains, oceanic lithosphere, associated intra-oceanic deposits, and the deformed margins of formerly independent Archean microcontinents (e.g., Bickford et al., 1990; Hoffman,
1988). In contrast, the Black Hills are dominated by rift-related metamorphosed
4 supracrustal rocks (Dahl et al., 2005a,b; Redden et al., 1990; Dutch and Nelson,
1990; Gosselin et al., 1988).
The broad objective of this research is to provide data and analysis to
further constrain the nature of the Early Proterozoic collision between the
Wyoming and Superior Archean cratons and the island arc terranes of the
Central Plains orogen.
Importance of the Mount Rushmore Quadrangle as a Proxy for Understanding the Tectonics of the Black Hills
The Black Hills were uplifted during the 75-35 Ma Laramide Orogeny and
expose a ~110 x 70 km portion of the buried Proterozoic collisional zone between the Wyoming and Superior provinces (Lisenbee, 1978; Bird, 1998). The
Precambrian rocks of the uplift record multiple deformational and metamorphic events related to the assembly of southern Laurentia (e.g., Redden et al., 1990;
Dahl et al., 1999b, 2005a,b). The Paleoproterozoic rocks have a complex sedimentary sequence and, in general, consist of basal metaconglomerates and quartzites nonconformably overlain by various quartzites, metagraywackes, iron formations, metavolcanics (tuffs, basalts), gabbros, mica schists, phyllites, and slates. The Precambrian rocks were affected by at least three deformational events (this study), although as many as five have been suggested (e.g. Redden et al., 1990), that resulted in multiple episodes of faulting and folding. Two episodes of prograde metamorphism and one regional-scale intrusive event are also recorded in the Paleoproterozoic rocks.
5 The Mount Rushmore quadrangle was chosen because of its complexity
and geologically critical location. The rocks of the quadrangle have undergone a
particularly complex history of Paleoproterozoic ductile folding, granite intrusion,
and metamorphism, including multiple deformational and metamorphic events.
Numerous faults, intrusions, and folding events are evident on existing large-
scale maps of the area (cf. Dewitt et al., 1989) and these attest to a complex
history of Proterozoic deformation. Three orogenic events have affected the
study area: (1) the ~1780 Ma Black Hills orogeny (Dahl and Frei, 1998; Goldich
et al., 1966), (2) the ~ 1720 Ma Dakotan orogeny (Chamberlain et al., 2002), and
(3) the ~1690 Ma Central Plains orogen (Sims and Peterman, 1986). The
earliest metamorphic event affected all Paleoproterozoic rocks of the Black Hills
and is temporally associated with the Black Hills orogeny. This event was later
overprinted by the thermal aureole of the Harney Peak intrusion at ~1715 Ma
(Norton and Redden, 1990; Dahl et al., 1999a). The Mt. Rushmore quadrangle
includes the northeastern contact of the Harney Peak Granite and a significant
portion of its metamorphic aureole. The Harney Peak granite intrusion played a
key role in syn-tectonic deformation and associated metamorphism in the Black
Hills. Understanding the metamorphic and fabric relationships in the southern
Black Hills is essential to developing a coherent tectonic model for the Black
Hills, which is, in turn, critical to understanding the assembly of southern
Laurentia.
The Mount Rushmore quadrangle is located approximately 25 kilometers
(15 miles) west-southwest of Rapid City and encompasses an area of roughly
6 225 km2 (Figure 3). It includes features formed during many of these events,
and metasedimentary, metavolcanic, and mafic igneous rocks, of the area record pre-, syn-, and post-intrusional fabrics and features that were reoriented along
the northeast contact of the Harney Peak Granite intrusion (Figure 4). The
Keystone
Figure 3 - Location and generalized geologic map of the Precambrian Black Hills. Modified from Redden and Lisenbee (1996). Black rectangle indicates approximate outline of the study area.
7
Figure 4 - Geologic map of the Mount Rushmore 7.5-minute quadrangle produced during this study. Geology is draped on the 1:24,000 scale DEM model for the quadrangle. Coordinates in UTM NAD 27. Note the outcrop pattern of the Harney Peak Granite (red) and the general north- northwest trend of major faults (bold black lines). See Plate I for more detailed geologic map.
8 incongruity of structural features and metamorphism related to the Harney Peak
Granite across fault-bounded blocks within the Mt. Rushmore quadrangle suggests a two-phase, protracted single orogenic event along the eastern margin of the Wyoming Province that encompasses both the Black Hills and Dakotan orogenies. Thus, the study area offers the best opportunity to unravel the complex stratigraphic, metamorphic, and structural relationships in the east- central Black Hills and elucidate the tectonic history of the Black Hills.
Research Objectives
The general objective of this study was the resolution of the metamorphic, stratigraphic, and structural relationships in the Mount Rushmore quadrangle, with the primary focus on the relationship between multiple deformational fabrics and multiple metamorphic events. The overarching goal of this work was to help elucidate the tectonic history of the Black Hills. The specific objectives of the research were to: (1) map the distribution of the major lithologic units; (2) provide information on stratigraphic relationships within the study area; (3) identify and quantify the geometry of major folding and faulting events; (4) identify and measure the orientation of the principal deformation fabrics and determine their relationships to each other, to metamorphic porphyroblast growth, and to regional structural, and tectonic events; and (5) assess the role of the Harney Peak
Granite in the deformational and metamorphic history of the region. These data are used to evaluate the degree to which correlations of structural features, deformation fabrics, and rock units can be made across multiple fault-bounded
9 blocks that make up the study area and to investigate the implications that such
correlations have for the overall tectonic development of the Black Hills. This
research provides data and analysis to assess geologic and tectonic
interpretations of the Proterozoic deformation in the Black Hills and the assembly of southern Laurentia.
Regional Geologic Setting
The Precambrian core of the Black Hills has been variously interpreted to
reveal a series of ensialic, epicratonic, and/or back-arc rift basins floored by
Archean basement rocks of the Wyoming (?) province infilled by Early
Proterozoic clastic sedimentary, volcaniclastic, and mafic igneous rocks (e.g.,
Gosselin et al., 1988; Redden et al., 1990; Redden and Dewitt, 1996). The
diverse tectonic interpretations demonstrate the imprecise and limited nature of
the current data base relating to the Black Hills. These rocks experienced at
least three episodes of penetrative deformation and two episodes of prograde
metamorphism from ~1780 – 1690 Ma. A late period of regional unroofing began
at ~1500 Ma (Holm et al., 1997). By the beginning of Cambrian time, the
Proterozoic rocks were unroofed and eroded close to their present structural
level (Lisenbee and Dewitt, 1993). During the Phanerozoic, the Paleoproterozoic
rocks were non-conformably overlain by a series of clastic and carbonate
sedimentary rocks (Lisenbee, 1978; Bird, 1998). During the Laramide orogeny,
area was subjected to large-scale doming that created the Black Hills and re-
exposed the core of the crystalline Precambrian complex.
10 Archean Basement
The oldest rocks in the Black Hills are Archean in age and consist of two small exposures, one at Bear Mountain (Bear Mountain Granite, 2563 ± 6 Ma;
McCombs et al., 2003), along the western edge of exposed Precambrian rocks, and the other (Little Elk Granite, 2559 ± 6 Ma; McCombs et al., 2003) along the
northeastern edge of the Precambrian core (Figure 3). The Little Elk Granite is a
medium- to coarse-grained, porphyroblastic, I-type granite gneiss that intruded a
supracrustal sequence of Archean biotite-feldspar gneiss (Gosselin et al., 1990).
The Archean biotite feldspar gneiss records a bedding parallel, east-northeast
trending, penetrative foliation that the intruded granitic gneiss does not, and thus
records the oldest deformation observed in the Black Hills (Gosselin et al., 1988).
Both the granite and the surrounding metasedimentary rocks have a dominant
northwest-trending foliation that crosscuts an older east-northeast-trending
foliation in the surrounding biotite gneisses (Gosselin et al., 1990; Redden et al.,
1990). The contact between Archean rocks of the Little Elk terrane and
Paleoproterozoic rocks has been interpreted as a fault (Gosselin et al., 1988) or
an angular unconformity (Jack Redden, pers. comm. in Gosselin et al., 1988).
The slightly older Bear Mountain Granite is a medium-grained to pegmatitic,
peraluminous, leucocratic, S-type granite that closely resembles the much
younger Harney Peak Granite (Gosselin et al., 1988, 1990; Redden et al., 1990).
The dominant structural element in the Bear Mountain Granite is a variably
developed northwest-trending foliation (Gosselin et al., 1990). The Bear
Mountain Granite occurs as concordant, massive to foliated, sill-like bodies that
11 were intruded into bedded quartz-feldspar-biotite schists (Redden et al., 1990).
The schists are overlain by a thin, discontinuous metaconglomerate that contains granite and feldspar pebbles, therefore the immediately enveloping quartz-biotite
schists are interpreted to be Archean while the metaconglomerate is interpreted
to be Proterozoic (Redden et al., 1990). Gosselin et al. (1988) report an east-
northeast trending, bedding parallel foliation in a couple of outcrops (Gosselin et
al., 1988; Redden et al., 1990), but it is not clear whether they occur in the
Archean rocks or overlying Paleoproterozoic rocks. Dahl et al. (2005a,b) report a
relict east-northeast trending fabric in garnets from a single sample (PR-1; Dahl
and Frei, 1998; Dahl et al., 2005a,b) of the Paleoproterozoic metasedimentary
rocks that surround Bear Mountain. From this, I infer that the east-northeast
trending fabric in the Bear Mountain terrane is restricted to the Paleoproterozoic
succession. Therefore, the east-northeast foliation observed at Bear Mountain is
younger than a similarly oriented penetrative deformation in the Archean rocks of
the Little Elk terrane. The fabrics represent two distinct penetrative deformational
events. The origin of the fabric observed in the Archean Little Elk terrane is
unknown while the fabric recorded in the Paleoproterozoic rocks surrounding the
Bear Mountain terrane has been related to cryptic tectonism between ca. 2170
and ca. 1960 Ma (Redden et al., 1990) or early stages of the Central Plains
orogen (Dahl et al., 2005a,b).
Metasedimentary rocks of Archean age do crop out within the
Precambrian core of the Black Hills (Gossellin et al., 1988; Redden et al., 1990).
Archean metasedimentary rocks are restricted to occurrences immediately
12 adjacent to Archean granites (Redden et al., 1990). Metasedimentary rocks of
known Archean age are practically indistinguishable from younger Proterozoic metasedimentary rocks, leaving open the possibility that unrecognized Archean sediments may occur elsewhere in the crystalline core of the Black Hills.
Paleoproterozoic Rocks
The Precambrian igneous and metamorphic rocks that are now exposed
in the Black Hills have been interpreted to be primarily of Paleoproterozic age
and have been divided into two broad age categories: (1) an older, ~2550 – 2480
Ma rift related sequence (Redden, 1981; Dahl et al., 2005a,b) that is unconformably overlain by (2) younger ~2015 – 1885 Ma (Bekker et al, 2003:
Dahl et al., 2005a,b) continental shelf-rise, shallow-water deposits (Redden et al.,
1990). At least two unconformities have been identified in the metasedimentary sequence (Redden et al., 1990). Most of the Paleoproterozoic metasedimentary rocks lack precise age constraints. In general, the Paleoproterozoic rocks consist of basal metaconglomerates and quartzites nonconformably overlain by various quartzites, metagraywackes, iron formations, metavolcanics (tuffs, basalts), gabbros, mica schists, phyllites, and slates.
~2550 – 2480 Ma Rift Succession
The oldest Proterozoic rocks are exposed in the Little Elk terrane and are described as basal quartz-pebble conglomerates of the Boxelder Creek
Formation (Redden, 1981; Gosselin et al., 1988; Redden et al., 1990). The
13 thickness of the Boxelder Creek was estimated to be >3300 m. The contact
between the Archean and Paleoproterozoic rocks in the Little Elk terrane has
been interpreted to be unconformable or faulted (Gosselin et al., 1988). The
Boxelder Creek Formation is interpreted to represent the older (pre-2480 Ma)
predominately non-marine, rift succession (Dahl et al., 2003, 2005a,b; Redden,
1981) and is restricted to the immediate area adjacent to the Little Elk Granite
and the Nemo area (Figure 3). It is described as a basal assemblage of chloritic
metaconglomerate, metamorphosed fanglomerate, taconite meta-conglomerate, and chloritic quartzite with pronounced lateral facies changes to phyllite and dolomitic marble overlain by fluvial metaconglomerate, quartzite, and oxide- facies iron-formation. An angular unconformity marks the top of this sequence
(Redden, 1981; Redden et al., 1990).
The oldest Proterozoic intrusive rocks (e.g., Blue Draw Metagabbro) have
a reported age of ~2480 Ma (Dahl et al., 2003) and predate the earliest
deformation recorded by the Paleoproterozoic metasedimentary rocks (Redden
et al., 1990). The Blue Draw Metagabbro occurs in the upper part of the
Boxelder Creek Formation, and therefore the age represents the minimum oldest
age of the Proterozoic sequence.
One period of folding predates the unconformity marking the top of the
older rift succession but after the emplacement of the Blue Draw Metagabbro,
bracketing a deformational event occurring in the older succession between
~2480 Ma and ~1985 Ma. This deformation is not recorded in the
Paleoproterozoic rocks of the Bear Mountain terrane. Dahl et al. (2003)
14 attributed this early deformation to the incipient breakup of the proposed supercontinent Kenorland.
~2015-1885 Ma Continental Slope-Rise Type Succession
The younger Paleoproterozoic succession nonconformably overlies the older rift-related succession (Redden et al., 1990; Dahl et al., 2003) and is more areally extensive. This younger succession dominates the exposed Precambrian metasedimentary rocks. Redden et al. (1990) subdivided the succession into two sequences based on inferred age relationships. The older sequence is composed of shallow-water clastic and carbonate rocks, overlain by tholeiitic metabasalts, slates, and phyllites (Redden et al., 1990). The younger sequence consists primarily of turbidite deposits, metagraywackes, shales, and volcaniclastic rocks in the central and southwestern Black Hills and metaconglomerates, debris flows, metabasalts, metavolcanics, iron formations, and quartzites in the east-central Black Hills (Redden et al., 1990).
The complicated faulting and folding, facies changes, spatial distribution and abundance of similar rock types across the Black Hills combined with the lack of precise age constraints and paucity of good stratigraphic marker units makes stratigraphic relationships across the entire Black Hills tenuous.
1715 Ma Harney Peak Granite
The youngest Precambrian rock in the Black Hills is the Harney Peak
Granite (Figure 3). The Harney Peak Granite intruded the younger continental
15 slope-rise succession at ~1715 ± 3 Ma (U-Pb monazite; Redden et al., 1990).
The Harney Peak Granite is a heterogeneous, peraluminous, leucogranite
consisting of an inner, biotite-rich, low-™18O granite, surrounded by a ring of
tourmaline-rich, high-™18O granite (Nabelek et al., 1992a). Walker et al. (1986)
conducted a Sm-Nd isotope study that demonstrated that the metasedimentary
rocks surrounding the Harney Peak Granite could not be its only source, and suggested a mixed Archean and Proterozoic source. Nabelek et al. (1992b)
determined that the variations in mineralogy and geochemistry of the Harney
Peak Granite are the consequence of source regions rather than post- emplacement differentiation. They concluded that the inner, biotite-rich granite
had an immature biotite-rich Archean metasedimentary source, whereas the
outer, tourmaline-rich granite was derived from structurally higher,
Paleoproterozoic muscovite-rich schists.
Deformation and Metamorphism
Three deformational events dominate the Precambrian Black Hills, although as many as five events have been suggested (e.g. Redden et al., 1990;
Holm et al., 1997; Dahl, 1999a). The relative timing of the deformational events is still under considerable debate (e.g., Dahl et al., 2005a,b; McCombs et al.,
2004; Chamberlain et al., 2002). The earliest deformation event is only recognized in the Archean Little Elk terrane and consists of a cryptic east- northeast foliation (Gosselin et al., 1988). This deformational event did not affect the Paleoproterozoic rocks in the Black Hills, and therefore it is inferred to be 16 Archean. The Little Elk Granite terrane also records the earliest metamorphic
events exposed in the Black Hills given that the oldest rocks are both gneisses.
The subsequent Proterozoic deformational events are the focus of this study.
Two metamorphic events are recorded in the Paleoproterozoic rocks of the
Black Hills, with the highest metamorphic grade reaching the second sillimanite zone southwest of the Harney Peak Granite (Figure 3; Helms and Labotka,
1991). After the deposition of the supracrustal sequences (ca. 1880 Ma; Dahl et al., 1999a), but prior to the intrusion of the Harney Peak Granite (~1715 Ma), the
Black Hills region underwent at least upper greenschist facies regional metamorphism (M1). The age of this metamorphic event has been broadly
interpreted to be between 1880 Ma and 1750 Ma (e.g., Redden and Dewitt,
1996; Redden et al., 1990; Dahl et al., 1999a, b). The M1 event has been
attributed to collision of the Wyoming and Superior provinces (Dewitt et al., 1986,
1989; Redden et al., 1990; Terry and Friberg, 1990; Dahl et al., 2005a,b). I has
been commonly reported in the olderliterature as having occurred at ca. 1840 Ma
based on a whole-rock Rb/Sr “disturbed isochron” date obtained by Zartman and
Stern (1967). However, reset whole-rock Rb/Sr data are highly suspect because
of limits of strontium rehomogenization (e.g. Lanphere et al., 1964; Herman et al.
1986; Faure, 1986). Recalculation of this data by Dahl et al. (1999a) yielded an
1800±70 Ma estimate, offering little improvement. Recent Pb-Pb dates of garnets and 40Ar/39Ar data from metamorphic rocks suggest 1770-1740 Ma
metamorphism for the regional M1 event (Dahl et al., 1999a, 1999b, 2000).
Although the more precise timing of this event has been unclear, recent chemical
17 dating of monazite by Dahl et al. (2005 a,b) suggests two episodes, the first at
~1780 Ma and the second beginning at 1755 Ma, based on a single metapelite
sample from the Bear Mountain area. The 1780 Ma age is similar to dates
reported in the Hartville uplift to the southwest of the Black Hills and has been
recently attributed to a “Black Hills orogeny,” with the younger ages ascribed to a
“Dakota orogeny” (Chamberlain et al., 2003).
The M2 event (~1715 Ma; U-Pb monazite age) is associated with the
emplacement of the Harney Peak Granite and hundreds of associated granitic
dikes and pegmatites (Norton and Redden, 1990). As such, the M2 event can be
thought of, with respect to the area of exposed Precambrian rocks in the Black
Hills, as a regional-contact event. The leucogranitic melts and pegmatites
associated with the M2 event are interpreted to have intruded relatively cool
o (<500 C) country rocks (Holm et al., 1997), overprinting the regional M1 metamorphism throughout the central crystalline core of the Black Hills. Nabelek et al. (in press) suggest that the Harney Peak Granite was intruded into a fault-
bounded crustal block that was subsequently uplifted to its current structural level. The regional M1, dominantly garnet-grade, assemblages were overprinted
by metamorphism and metasomatism that accompanied intrusion of the Harney
Peak Granite.
Tectonic Models
Two different tectonic models have been proposed for late Archean orogenesis along the southeastern margin of the Wyoming Province, to account
18 for both the presence of the Little Elk and Bear Mountain granites and their
weakly developed foliations. Karlstrom and Houston (1984) proposed that an
unidentified Archean continental block collided with the Wyoming Province,
resulting in a continent-continent collision and creating a coherent continental
mass that was later rifted to form the basin(s) into which the majority of the Late
Proterozoic metasediments in the Black Hills were deposited. In contrast,
Redden et al. (1990) suggested that two episodes of ensialic rifting generated and exposed the Little Elk and Bear Mountain terranes prior to 1964 Ma and led to the development of the Black Hills basin(s). Currently available data and observations are not sufficient to favor either model.
Models of Paleoproterozoic tectonism in the Black Hills are no less controversial. Redden et al. (1990) suggested a tectonic history that included, in order of occurrence: an early compressional event of undetermined orientation, ensialic rifting, NE-SW compression, ensialic rifting, subsidence, E-W compression, subsidence, NW-SE compression (F1 folding), E-W compression
(D2 event), intrusion of the Harney Peak, and finally NW-SE compression.
Klasner and King (1990) developed a tectonic model based on the assumption
that the Black Hills are the southern extension of the Trans-Hudson Orogen
(THO). New age data (cf. Dahl et al., 1999 a,b; Chamberlain et al., 2002), which
are discussed in a subsequent section of this report, suggest that the intervening
orogen between the Wyoming and Superior provinces (referred to as the Dakota
segment) was formed by a separate orogenic event and therefore calls into
question the Klasner and King model. I propose an alternative tectonic model
19 based on the lack of structural continuity and stratigraphic correlation across the
fault boundaries in the study area.
Unresolved Issues of Black Hills Geology
The preceding overview of the regional geology suggests that there are
several unresolved questions related to the construction of the Precambrian
Black Hills that bear on the broader tectonic history. In particular, (a) it is uncertain that the Black Hills represent a simple collisional terrane of the Trans-
Hudson orogeny, and (b) the recent propositions that closely-spaced multiple deformations in time represent multiple orogenies is not consistent with collisional orogens, such as the Himalayas, that are constructed over tens of millions of years. Moreover, the data presented in this study suggest that fault- bounded blocks within the Black Hills may have had different structural, stratigraphic, and tectonic histories. New research (e.g., Nabelek et al., in press;
this study) is questioning the view that all the rocks of the Precambrian Black
Hills were subjected to the same deformational and metamorphic events. Many
questions regarding the nature of deformation and the causative events in the
Black Hills remain unanswered in sufficient detail. Among the broader questions
are: What tectonic event(s) are responsible for the multiple tectonic fabrics
observed in the Black Hills? What timing constraints (if any) can be applied to the observed metamorphism and deformational fabrics? What role did the
Harney Peak Granite play in regional deformation and metamorphism, or
20 alternatively, what role did deformation play in emplacement of the granite magmas? Are there unrecognized terrane boundaries in the Black Hills?
Lithotectonic Units versus Stratigraphy
The study area is a prime example of repetition of similar rock units that have complicated structural, metamorphic, and facies relationships that are not easily classified into any determinate stratigraphy. In general, the dominant rock assemblages of Proterozoic age in the study area are (1) pelitic schists, phyllites, and slates; (2) metagraywackes, siliceous mica-schists, and quartzites; (3) metabasites – including basalt, gabbro, and amphibolite; and (4) granite and associated pegmatites. Minor rock assemblages include iron-formations, calcareous phyllites, and calc-silicates.
Structural repetition by folding and faulting, and facies changes within the supracrustal units, make stratigraphic correlations tenuous at best. For example, the Mount Rushmore quadrangle is cut by at least four major north-west trending faults that divide it into nine fault-bounded blocks (Figure 5), each with a potentially different structural history. In general, the fault-blocks contain similar lithologic units but precise correlation between them remains problematic, hence no definitive stratigraphic correlation between rocks of separate fault blocks has been established.
This study relies on grouping rocks into mappable lithostratigraphic sequences within the fault-bounded blocks. The lack of distinct marker units
21 within the metasedimentary sequences, along with varying degrees of metamorphic overprinting and facies changes, and lack of any age dating techniques applied to the metasedimentary sequence made determination of an actual stratigraphy nearly impossible. Where possible, stratigraphic and structural facing indicators were used to determine stratigraphic younging directions. Correlation with other workers’ stratigraphic successions would be, at this time, speculation and has been avoided (cf. Appendix B)
The rocks of the study area have been divided into fourteen lithostratigraphic units. The Harney Peak Granite and related pegmatites are clearly the youngest Precambrian rocks cropping out. Age relationships of other
Paleoproterozoic rocks remain equivocal. Similar rock types occur in most of the fault-bounded blocks and have been mapped as similar rock units. This grouping is not intended to imply any stratigraphic correlation between similar lithologies across major faults within the study area. Appendix A lists and describes the lithostratigraphic units used in this study (Table 1). Plate I is the geologic map produced based on these units.
22
Figure 5 - Outline map of fault bounded blocks delineated during this study in the Mt. Rushmore 7.5-minute quadrangle. Letter designations for fault-blocks will be used throughout this paper. A major structural boundary exists along the northeast side of fault-block C (Keystone Fault) within the study area.
23 Harney Peak Granite and Mainly qtz-schist, qtz-mica schist +/- grnt Iron formation and Fe-rich schists; similar to but related pegmatites and/or stuarolite schist with interlayered slates not quite true BIF; alternating layers of cherty and metasiltstone. Minor metagraywacke quartzite with Fe-rich layers
Qtz-mica schists +/- garnet away Grey to black slates and "spotted" slates/schists from granite; sillimanite with subordinate metasiltstone Quartzite and cherty quartzite near granite
Turbidite sequences: Metagraywacke to clean Qtz-mica +/- garnet and/or Interlayered metagraywacke and slates with quartz schists with late staur and grnt staurolite "spotted" schists minor schistose layers porphyroblasts. Distinctive felsic, "leucocratic"
Qtz-biotite-grnt schists, graphitic Arkosic metagraywackes with interlayered qtz-sch in part; Qtz-mica-grnt-st schists with qtz Qtz-schists and qtz-mica schists Some recognizable turbidites. More micaceous an pebbles (matrix supported metaconglomerate) "melanocratic" than above, but similar.
Qtz-mica schists and metagraywacke; Qtz phyllite and phyllonitic qtz schist; Qtz-biotite-grnt schist and metagraywacke; graphi some graphitic schists interlayered; lower grade unit same(?) as above in part similar to above
Massive to thick bedded metagraywacke Dark graphite and sulfide rich schists and shales; with interlayered qtz schist +/- grnt Biotite schist +/- grnt highly altered in many places with residual "coke" and/or staur may contain retrograde late garnets
Staurolite schist; qtz-mica schist with large Amphibolite, amphibolite schist, metagabbro, staurolite porphyroblasts; distinctive unit and metabasalt undifferentiated
Table 1 - Summary of lithostratigraphic units used in this study. Colors correspond to the map shown in Figure 4. See Plate I and Appendix A for more detailed description.
Problems with Stratigraphic Correlation
One objective of this study was to provide information on stratigraphic relationships within the study area. Many previous structural interpretations in the Black Hills have been based on stratigraphic relationships (Noble and
Harder, 1948; Redden 1963, 1968; Ratte and Wayland, 1969; Bayley, 1972b,
1972c; Ratte, 1986), so a significant effort was made to verify stratigraphic interpretations that could potentially be used to assist the structural interpretations of this study. However, for the reasons stated above, stratigraphic correlations are difficult. Available age data are primarily derived from metamorphic porphyroblasts (e.g. garnet, staurolite, monazite) (e.g. Gardner et
24 al., 1996; Holm et al., 1997; Dahl et al., 1999a, b; Dahl et al., 2003, 2005a,b) that
give no information about stratigraphic successions.
In detail, the difficulties in defining and correlating stratigraphic units with
any confidence within the crystalline Black Hills are three-fold. The first, and
most basic problem, is the recurrence of identical rock types within the
succession and the paucity of good stratigraphic marker units. Second, the rocks
within the crystalline core have been multiply and complexly deformed. As a
result, detailed structural analysis is required to unravel the stratigraphic
relationships. Thirdly, the thermal overprint of the Harney Peak Granite has
significantly metamorphosed the country rocks within its aureole and the degree
of metamorphism is not uniform (Nabelek et al., in press).
Locally, and sometimes for some distance in the field, stratigraphic
relationships may be determined based on primary structures and distinctive
sequences. It is very difficult, however, to correlate these sequences over
significant distances due to the structural and metamorphic complexity of the
study area and the general similarity of lithostratigraphic units. For example, a turbidite succession at point A may look very similar to another turbidite succession at point B. However, the similarity does not necessarily imply a stratigraphic continuity between the two turbidite successions. Only careful and consistent observation between points A and B will yield a correct interpretation, and only if there is sufficient outcrop to gather the necessary data.
By far, the most consistent primary structure in the study area that provides stratigraphic topping directions is graded bedding, but incomplete
25 exposure and folding yields reliable younging interpretations only very locally. A
large portion of the Mt. Rushmore quadrangle includes turbidite successions that
locally preserve incomplete Bouma sequences. Primary bedding features may
be recognized in some layers, but the inconsistent nature of preservation of such
features between outcrops makes any interpretation of structure based solely on stratigraphic younging criteria suspect.
Previous Stratigraphic Correlations
The stratigraphy of the Paleoproterozoic medisedimentary rocks in the
Black Hills remains nebulous. Redden (1963, 1968) reported stratigraphic
relationships for metasedimentary rocks in the Fourmile and Berne 7.5-minute
quadrangles southwest of the study area. Redden et al. (1990) proffered a
rough, first-order correlation of stratigraphic relationships in the Precambrian
rocks across the entire Black Hills. However, many of the assertions and
correlations of Redden et al. (1990) remain unclear and subject to debate.
Previous workers raised concerns about correlating rocks across multiple faults
based solely on lithologic similarity (e.g. Ratte and Wayland, 1969; Norton, 1974;
Woodland, 1979; Ratte, 1986). Appendix B summarizes some of the major
inconsistencies with previous stratigraphic correlations and their application to
ascertaining the regional structure.
26 Methodology
This study was implemented in two phases: (1) a field component and (2) a laboratory component. Detailed mapping began in late May, 2000, and continued through mid-August, 2000. Additional field seasons corresponding to the same time frame were undertaken in 2001 and 2002. Laboratory work was conducted during the intervals between field seasons and continued after the completion of the mapping phase.
Mapping was done at 1:12,000-scale or smaller and compiled at the
1:24,000-scale using USGS digital topographic base maps. Orientation measurements of foliations, fold axial planes and hinge lines, mineral and intersection lineations, and fractures were collected using a BruntonTM Model
5005LM Pocket Transit. Locations were recorded in UTM coordinates (1927
North American datum) using a Garmin GPSmap76 hand-held GPS unit equipped with WAAS-differential correction (Wide Area Augmentation System).
Typical positional accuracy was better than three meters. Field work involved traverses, generally across regional strike, along all major streams, ridges, roads, and railroads. Additional area-specific traverses were undertaken and spaced closely enough to permit tracing of particular contacts. In addition, “float” was identified and plotted on the map to help delimit contacts.
Structural analyses were aided using the stereographic projection software: GEOrient developed by Professor Rod Holcombe (Department of
Earth Sciences, University of Queensland, Australia) and SpheriStat™ version
2.2 from Pangaea Scientific. Laboratory studies involved petrographic and
27 micro-structural analysis using an Olympus BX40 polarizing microscope. Two
hundred and twenty-nine oriented rock samples were collected; 87 of these
samples were cut to make thin sections. Fifty-one additional thin sections from
the study area were provided by Dr. Peter Nabelek and Dr. Robert Bauer.
Previous Geologic Mapping
The earliest recorded geologic investigation was conducted and led by
General W. F. Harney in 1855 at the bequest of the United States Army
(Froiland, 1990). In 1859, Captain W.F. Raynolds and Dr. F. V. Hayden led
another expedition into the Black Hills and published the first geological report of
the Black Hills in 1869. General George A. Custer led an expedition into the
Black Hills in 1874, with Professor N.H. Winchell as geologist, which discovered
the presence of gold in the Black Hills. As a direct result of the discovery of gold,
the Bureau of Indian Affairs ordered a survey of the Black Hills in 1875, which
was directed by Henry Newton and Walter P. Jenny, with fifteen assistants and
an escort of 400 soldiers. The results of the survey were published in 1880.
One of the earliest studies conducted in the crystalline rocks of the Black
Hills was conducted by Van Hise (1890) who noted the vertical N-S trending slaty
cleavage and folds in the vicinity of the Harney Peak Granite. Darton and Paige
(1925) first mapped this area as part of the Central Black Hills Folio. Hess
(1925) and Landes (1928) published the first reports on pegmatites in the study area. Connolly and O’Harra (1929) first described the geology of ore deposits.
Balk (1931) studied the metamorphic enclaves and foliations of the Harney Peak
28 Granite. During World War II and after, the Black Hills were a locus of numerous
investigations by the United States Geologic Survey (e.g., Fisher, 1942, 1945;
Gwynne, 1944; Noble and Harder, 1948; Noble et al., 1949; Sheridan et al.,
1957; Norton et al., 1962). Previous mapping of the study area was conducted
by Norton (1976). Mapping adjacent to and within close proximity of the study
area was conducted by Redden (1963, 1968), Ratte and Wayland (1969), Bayley
(1970, 1972a, 1972b), Woodland (1979). Dewitt et al. (1986) published a
1:250,000-scale map of the Black Hills.
STRUCTURAL RELATIONSHIPS AND DEFORMATIONAL HISTORY
The Black Hills have undergone multiple episodes of both prograde
metamorphism and deformation. The rocks in the study area are primarily of
mid- to upper greenschist metamorphic facies except where they have been
affected by the metamorphic aureole of the Harney Peak Granite (Figure 6). In the southwestern portion of the study area, rocks immediately adjacent to the
Harney Peak Granite contain flattened aggregates of sillimanite with the highest
zone recording coexisting sillimanite + garnet + biotite (cf. Helms and Labotka,
1991). The highest grade metamorphic rocks in the Black Hills occur southwest
of the study area, where the second sillimanite isograd occurs adjacent to the
Harney Peak Granite (Figure 6; Duke et al., 1988). The timing of Proterozoic
deformational and metamorphic events is constrained between 1850 Ma and
1555 Ma (Terry and Friberg, 1990; Redden and Dewitt, 1996). Chamberlain et
29 al. (2002) proposed two periods of east-west collision: (1) the ~1778 Ma Black
Hills Orogeny resulting from the collision of the Dakota block with the Wyoming province and (2) the ~1720 Ma Dakotan Orogeny as the terminal collision of the
Wyoming and Superior provinces. Major episodes of folding and faulting culminated with the intrusion of the Harney Peak Granite at 1715 Ma (207Pb-206Pb age, Redden et al., 1990), although at least one penetrative fabric within the study area apparently post-dates the Harney Peak Granite (this study).
Summary of Previous Structural Studies
The earliest deformational fabric in the Black Hills occurs in the Archean
Little Elk terrane and is restricted to the Archean rocks of the terrane (Gosselin, et al., 1988). The Little Elk terrane has two distinct fabrics: an older east- northeast-trending foliation found in the surrounding (Archean) biotite-feldspar gneisses and a younger, strong northwest-trending foliation in the granite proper that cross-cuts the earlier northeast-trending fabric (Gosselin et al., 1990).
Gosselin et al. (1990) interpreted the younger foliation to be contemporaneous with the emplacement of the Little Elk Granite and suggested, based on the 1850
Ma Rb-Sr whole rock age of Zartman and Stern (1967), and the similar orientation of the dominant fabric in Proterozoic rocks elsewhere in the Black
Hills, that this foliation was reactivated during Paleoproterozoic deformational events. The Bear Mountain terrane also records an earlier, weak east-northeast trending bedding-plane foliation in the encapsulating younger sequence of
30 Paleoproterozoic metasedimentary rocks and a younger concentric foliation that outlines the structural dome seen both in the Archean schists and, to a lesser degree, in the Bear Mountain Granite sensu stricto (Gosselin et al., 1990;
Redden et al., 1990). This younger foliation is interpreted to have formed contemporaneously with doming that is postulated to be a result of the emplacement of an underlying unexposed pluton of the Harney Peak Granite based on 1680±25 Rb-Sr muscovite ages from the flank of the Bear Mountain dome (Zartman, in Ratte, 1986; Redden et al., 1990). Dahl et al. (2005a,b a,b) reported an early ENE-trending fabric preserved in the cores of garnet porphyroblasts from the Bear Mountain dome. The east-northeast foliation observed in the younger Paleoproterozoic sequence overlapping the Bear
Mountain terrane has been related to early Central Plains orogeny ca. 1775 Ma
(Dahl et al., 2005 a,b).
The most complete summary of deformational events affecting the
Proterozoic rocks of the Black Hills is given in Redden et al. (1990). The earliest event, D1, has been inferred from the cryptic presence of east-northeast-trending, north-vergent, F1 fold nappes recognized on the basis of top and bottom criteria
(Dewitt et al., 1989; Redden et al., 1990). No penetrative fabric or metamorphism are associated with the D1 event in the Black Hills, although the early ENE-trending fabric reported by Dahl et al. (2005 a,b) may indicate that this event was locally penetrative in nature. The timing of this event is unclear, but it may represent the earliest arrival of a southward-younging (1790-1630 Ma)
31
Figure 6 - Metamorphic zones of the Precambrian Black Hills (from Nabelek et al., in press). Red box is approximate location of study area. BM = Bear Mtn granite; HPG = Harney Peak granite; LEC = Little Elk Granite.
32
Figure 7 - Generalized geologic map of the Precambrian Black Hills showing location of the Archean terranes and major structural features. Note the overall pattern of faulting and resultant displacement of mappable units northeast of the Harney Peak Granite. Mapped adapted from Dewitt et al. (1986), Ratte and Wayland (1969), Ratte (1986), and unpublished maps of the Silver City quadrangle. White box represents approximate location of the study area.
33 collage of island-arcs that make up the Central Plains orogeny (Dewitt et al.,
1989; Redden et al., 1990; Dahl et al., 1999a; Sims and Peterman, 1986;
Bickford et al., 1986; Sims et al., 1991). Dahl et al. (2005 a,b) suggest a 1775 ±
10 Ma age for this event.
The dominant structural fabric in the Black Hills is associated with the D2 deformational event and associated regional metamorphism (M1). The D2 event
produced north- to northwest-trending upright folds (F2) that refolded F1 folds
(Redden et al., 1990). The D2 event is generally considered to be associated
with east-west collision of the Wyoming and Superior provinces. D2 produced a
locally penetrative, steeply dipping axial-planar foliation (Dewitt et al., 1989).
This axial-planar foliation has been designated S2 because an earlier fabric (S1)
is preserved as oblique inclusion trails in regionally metamorphosed garnets
(Dahl et al., 1999a). However, it is not clear whether S1 is associated with D1,
represents the early stages of D2, or represents garnet overgrowth of a primary
compositional banding/bedding unrelated to deformation (i.e. is actually S0).
The cryptic D3 event resulted in F3 “cross-folds.” D3 is of uncertain age
and origin, but purportedly refolds F1 and F2 folds in the northern and central
Black Hills, most notably in the Pactola Reservoir area just north of the study
area and near the Dearfield area east of Silver City (Redden et al., 1990). The
relationship between D3 and other fabrics remains uncertain. F3 folds have been interpreted to be a late-stage modification of F2 folds by a major shear couple
affecting the Black Hills during the waning stages of continental assembly
(Redden et al., 1990; Dahl et al., 1999a). Localized doming of country rocks (D4)
34 and metamorphism (M2) accompanied emplacement of the Harney Peak Granite and hundreds of associated granitic dikes and pegmatites (Norton and Redden,
1990). D4 produced local folds and schistosity associated with doming, locally
transposing the dominant S2 fabric into S4. A post intrusion D5 event, which produced a northeast-trending fabric across the central Black Hills ca. 1535 Ma, has been suggested by Redden et al. (1990), presumably as a far-field response to the Central Plains orogen. Table 2 is a summary of deformation and related structures as outlined above.
Despite the structural framework discussed above and outlined in Table 2, many questions regarding the structural history of the Proterozoic Black Hills remain unresolved. Previous workers have assumed that all the
Paleoproterozoic rocks of the Black Hills have experienced the same sedimentologic, tectonic, and structural histories and have constructed a complicated tectonic model to explain this assumption (e.g., Redden et al., 1990;
Terry and Friberg, 1990; Helms and Labotka, 1991; Dahl et al., 1999 a,b, 2005 a, b). For example, the oldest Paleoproterozoic rocks are restricted to a few fault- bounded blocks in the Nemo area and do not outcrop elsewhere in the Black
Hills. Should not a vestige of this sequence be seen in the temporally equivalent
Bear Mountain terrane? The lithologic differences between parts of the Black
Hills separated by faults has been attributed to multiple ensialic rift basins
(Redden et al., 1990) which leads to an abstruse explanation of Black Hills
35 Folds Fabrics Style Orientation S – relict foliation in Northeast-trending, 1 garnet representing early Tight to Isoclinal overturned to north- D1 stages of D (Dahl et al., northwest 2 1998a)
S – steeply dipping axial North to northwest 2 Isoclinal planar foliation D2 trending, upright
*S1 of Redden et al., (1990)
Upright, tight – More westerly than F2 isoclinal, doubly folds to east-west, S – generally non- plunging (Redden, 3 D3 moderate to steep penetrative foliation (?) pers. comm., 2001); plunges deforms F2
Doming and S4 – schistosity D4 associated local folds; Variably oriented associated with granite refolds earlier folds emplacement
S5 – northeast trending axial-planar foliation to Small, late folds Northeast trending spaced cleavage D5 locally recognized
*S4 of Redden et al., (1990)
Table 2 - Summary of Proterozoic deformational events and related structures. See text for explanation. Data summarized from Redden et al., (1990), Dahl et al., (1998), Dewitt et al., (1989), and Noble et al., (1949).
tectonism. This follows from the assumption that all the rocks have the same histories, which is likely not the case. Moreover, a multitude of deformational events have been ascribed to the Black Hills but not all parts of the Black Hills record all episodes of deformation. The following section discusses the
36 relationship between multiple deformational fabrics within the Mount Rushmore
quadrangle and their relationship to previous studies.
Deformational Events and Features Considered in this Study
The study area has been divided up into nine fault-bounded blocks (Figure
5). These fault-blocks have different structural histories and are used as
structural domains in this study. The attitude of structural features is first
discussed with respect to the major deformational events (Table 2) for the study
area as a whole. The variations among the structural domains are then discussed to help elucidate the structural history of the study area.
Three distinct deformational events have been recognized in the study area. These events correspond to the D2, D4, and D5 regional events as outlined
in Table 2. I have chosen to refer to these events within the regional context of
deformational events reported for the Black Hills (e.g. D2-S2-F2, etc.) despite the
fact that within the study area and fault-blocks they represent local D1, D2, and D3
events and related structures. Unnecessary complication will be avoided by
referring to these events with respect to the overall regional context as reported
in the literature and will allow the reader to evaluate this study more easily in the
regional context. However, it must be recognized that in doing so, this author is not implying that the deformational history as outlined in Table 2 is strictly applicable to the study area. Moreover, not all fault blocks record all deformational events.
37 Bedding, Primary Sedimentary Structures, and Compositional Variation
Much of the Mount Rushmore quadrangle is underlain by moderately metamorphosed sedimentary sequences, which in many instances preserve primary sedimentary structures. Bedding (S0) is the most prevalent primary structure recognized, but is not ubiquitous, nor consistent, throughout the study area. Where it is recognized unequivocally, bedding consists mainly of thinly laminated compositional variations in finer grained metasedimentary rocks whereas thicker beds preserve primary sedimentary structures such as cross bedding, particularly in coarser metagraywackes. Graded bedding characteristic of turbidites is common. Disrupted bedding and soft-sediment deformation are also recognized.
Most metasedimentary rocks show some compositional banding that has been interpreted as bedding by previous workers, but is not likely to be primary layering formed during deposition (Figure 8). As such, it cannot be bedding sensu stricto and is referred to herein as compositional banding (Sc). In many cases, true bedding and compositional banding may be present simultaneously.
Where this occurs, bedding (S0) is usually sub-parallel to the compositional banding (Figure 9). In other instances, compositional banding may look essentially like bedding at the meso-scale and can only be correctly identified microscopically (Figure 10). The origin of the compositional banding may be a compaction
38 A.
B. Figure 8 - (A) Typical lichen covered exposure of metasedimentary rocks within the study area. Note the prominent layering of the rocks and the younger spaced cleavage trending from lower right to upper left in the left exposure. Most previous workers have generally designated layering such as this as bedding (S0). But is it? The difficulty is not recognizing the layering but determining the origin of the layering. See Figure 9 for comparison. (B) Overlay showing younger spaced cleavage direction (red lines). Hammer for scale is 32.5 cm long.
39 A.
B.
Figure 9 - Atypical exposures of the typical relationship of bedding (S0) to compositional banding (Sc) in metapsammites within the study area. (A) Sc trends gently downward from left to right. Note the abrupt change in dominant lithology and grain-size in the upper part of the photograph delineating Sc. Bedding (S0) is subparallel to Sc and is expressed as mm-scale banding in the center of the photograph. The blue pencil is parallel to small shear bands. Fortunately, this small section of the outcrop was not completely covered in lichen permitting the recognition of S0. Pencil is 14 cm long. (B) Sheared S0 (red arrows) is cut by dominant S2 fabric, which trends at a shallow angle from lower left to upper right. Note US dime for scale.
40 A.
B. C. Figure 10 - Example of bedding recognition through petrographic analysis. (A) Plane polarized light scan of representative thin section MR266, a quartz-mica schist. Note the overall layered appearance of the section and the younger spaced cleavage trending subvertically (red arrow) bottom-right to upper-left that is especially prominent in the center right of the photomicrograph. Compositional banding (Sc) may be easily mistaken for true bedding in the field. Yellow box located left center is approximate field of view for (B) and (C). Thin section is standard size 27 mm X 46 mm. (B) Magnified view of area delineated by yellow box in part (A). Note true bedding in center of photomicrograph trending top left to bottom right between zones of dominant foliation (Sc). Bedding is apparently sheared along bottom contact. Field of view is approximately 0.3 by 0.5 mm. (C) Same as (B) with overlay showing bedding S0 (yellow lines) and dominant foliation Sc (blue lines) for clarity.
41 foliation, tectonic in origin or perhaps an effect of metamorphic segregation.
Regardless, great care was taken during this study when assigning the term
“bedding” because of the structural and tectonic implications implied by its usage.
The use of the term bedding was restricted to areas where I was convinced that the structure was primary, for example recognition of graded bedding or cross- beds. In all other instances, I refer to the foliation as compositional banding (cf.
Figures 8 – 10).
D1 Deformation
Regional D1 deformation is attributed to a series of north-directed nappes emplaced as thin-skinned foreland thrust complexes that overrode the southeastern Wyoming craton (Dewitt et al., 1989; Redden et al., 1990; Dahl et al., 1999a, 2005a,b). The Cheyenne fold belt has been interpreted as the root of these nappes (Condie and Shadel, 1984; Duebendorfer and Houston, 1987;
Premo and Van Schmus, 1989; Houston et al., 1989; Dahl et al., 1999a). The
Black Hills would represent the northern extent of nappe emplacement during the onset of the 1780-1740 Ma Medicine Bow orogeny (Chamberlain, 1998).
However, Redden et al. (1990) constrained the age of D1 deformation to 1884-
1840 Ma, and Dahl et al. (1999a, 2005a,b) suggested a circa 1790-1780 Ma age for onset of D1 based on biotite cooling ages and U-Th-Pb monazite ages.
Chamberlain et al. (2002) proposed a ~1778 Ma date for this early event in the
Laramie Range based on U-Pb ages of syn-deformational sphene. Recent work
42 in the Black Hills by Dahl et al. (2005a,b) has constrained the D1 event between
1785 and 1775 Ma. The attribution of the D1 event remains uncertain.
Foliation
Previous workers have reported that no fabric is associated with D1 (e.g.,
Redden et al., 1990), however Dahl et al. (2005 a,b) report a presumed S1 relict fabric in garnet porphyroblasts from a single sample taken from the younger
Paleoproterozoic sequence outcropping ~1 km east of Bear Mountain. Evidence for an S1 foliation was not observed in the study area, with the possible exception
of fault-block B. Rocks in fault-block B typically have two fabrics, original
bedding and a weak foliation that is axial planar to recumbent tight to isoclinal
folds that have shallow axial planar dips (Figure 11). This axial planar fabric may be related to D1, although that interpretation is equivocal. Two possibilities exist
for the paucity of early S1 fabrics: (1) the penetrative nature of later
deformational events throughout the majority of the study area transposed S1, or
(2) the fault-bounded blocks along the east side of the Black Hills do not record this earlier event. The absence of pre-Harney Peak porphyroblasts in the study area as compared to areas west of the Keystone fault supports the latter possibility but does not preclude the former. It is also possible that the absence of pre-Harney Peak porphyroblasts could be because the rocks never reached sufficient temperatures to produce M1 related porphyroblasts. The origin of the
compositional banding (Sc) described above may also be related to D1
deformation.
43 Dahl et al. (1999a) reported inclusion trails in garnets that are oblique to
the S2 foliation and may be relict fabrics associated with D1. Dahl et al. (2005a,b)
reported relict east-northeast-trending foliations in garnet porphyroblasts in two
samples of metapelite from the Bear Mountain dome. However, it is unclear how
the recognition of an S1 fabric in a single sample by Dahl et al. (2005a,b) in
kyanite grade rocks at the Bear Mountain locality (i.e., presumably deeper crust)
may be evaluated with respect to the lack of penetrative S1 fabrics reported by
earlier workers (e.g. Redden et al., 1990). Dahl et al. (1999a; 2005a,b)
presented no argument to preclude their reported S1 from being preserved
bedding. Weak axial planar foliations of early formed recumbent isoclinal folds
observed west of the study area in the Hill City quadrangle, similar to that
observed in fault-block B, may also be S1. However, as discussed below, it is likely that fault-block B has a distinct structural history and that the weak axial planar fabric is related to the D2 event.
Folds
Unequivocal evidence for the F1 folding was not recognized in the Mt.
Rushmore quadrangle. Two examples of large-scale F1 folding that occur in the
study area were shown to this author by Dr. Jack Redden of the South Dakota
School of Mines and Technology (retired). Both examples are described briefly
in Redden and Duke (1996), stops one and two, in the road log of field trip nine
of the field guide. In both instances, I remain unconvinced of the interpretation of
the outcrops. The first example, located in the extreme northeastern portion of
44 the Mt. Rushmore quadrangle, is based on the recognition of a northeast-
trending “antiformal syncline” as determined by graded-bedding on opposite
limbs of the fold. After multiple, careful examinations of the outcrop, I have not
been able to convince myself of the unequivocal existence of the fold, nor did I
find any evidence for such folds in the immediate environs. The second example
is located along the central eastern-edge of the Mt. Rushmore quadrangle. In
this case, the main argument for the existence of early F1 folds is based on
lineations that wrap across perceived fold hinge-lines (Redden, 2001 pers.
comm.). Even if the early folds exist, wrapping of early-formed lineations across
the hinge-lines of the folds can be easily explained by progressive noncoaxial
strain during folding (Hobbs et al., 1976, pg. 192).
Although little evidence for large-scale F1 folds exists in the Mt. Rushmore
quadrangle, bedding (S0) has been locally reoriented by F1/F2(?) folds. However,
as discussed below, these folds may be F2 folds that have been locally
reoriented by the emplacement of the Harney Peak Granite. Occurrence of
F1/F2(?) folds in the study area is ostensibly restricted to fault-blocks A and B
(Figure 5), but F1 folds have been reported by other workers in the Mount
Rushmore quadrangle and throughout the Black Hills (Redden and Duke, 1996;
Redden, 2001 pers. comm.). Where recognized by other workers, F1 folds have
generally northeast-trending orientations and are moderately to steeply plunging.
Recognition of early F1 folds in the Mount Rushmore quadrangle is difficult because of transposition of S0 by later fabrics. However, within fault-blocks A
and B (Figure 5) north of the Harney Peak Granite, these early F1/F2(?)
45 A.
B.
Figure 11 - (A) Photograph of early recumbent fold of interbedded metagraywacke and micaceous schist east of Hill City on Highway 16/385, just west of the study area in the Hill City quadrangle, looking N. A weak axial planar cleavage (S1), roughly parallel to the road, is evident in this early F1-fold. Multiple folds like this exist in fault-block A of the study area, but are not as well exposed. Note the 1987 Chevy Blazer for scale. (B) Weak axial-planar fabric of the early F2 folds shown in hand-sample. “Spots” trending across the layer from lower left to upper right define the foliation and are typically biotite clots that surround quartz centers. Layering is likely S0 and the spots define local S1. Layer is approximately 8 cm thick.
46 structures are generally tight to isoclinal, recumbent folds with gently dipping
axial surfaces that either lack or have a very weak axial planar fabric. These
folds are similar to folds in the Hill City quadrangle, west of the Mount Rushmore
quadrangle, that may be related to D1 (Figure 11).
D2 Deformation
The dominant structural features and M1 regional metamorphism in the study area are associated with the D2 event. D2 is attributed to the east-west collision of the Wyoming and Superior provinces and has been widely interpreted to be the southern extension of the Trans-Hudson orogen (Green et al., 1979,
1985; Sims and Peterman, 1986; Van Schmus et al., 1987; Klasner and King,
1990; Houston, 1993; Baird et al., 1996), although strict usage of the term Trans-
Hudson is problematic. For instance, Nabelek et al. (2001) and Chamberlain et al. (2002) suggested that the D2 event was a result of east-west collision between the Wyoming province and the Dakota block, an intervening crustal block of uncertain origin recognized from geophysical data (e.g. Baird et al.,
1996). Redden et al. (1990) constrained the timing of the D2 event between
1880 Ma and 1715 Ma, prior to the intrusion of the Harney Peak Granite (1715
Ma) and after the deposition of the supracrustal sequences (ca. 1880 Ma; Dahl et
al., 1999a). Redden et al. (1990) assigned an 1840±70 Ma age for D2
deformation based on Rb-Sr whole rock “disturbed isochron” data from the
Archean Little Elk Granite reported by Zartman and Stern (1967). However,
reset whole-rock Rb/Sr data are highly suspect because of limits to Sr
47 rehomogenation (eg. Lanphere et al., 1964; Herman et al., 1986; Faure, 1986).
Recalculation of these data by Dahl et al. (1999a) yielded an 1800±70 Ma
estimate, offering little improvement. Recent Pb-Pb dates of garnets and
40Ar/39Ar data from metamorphic rocks suggest 1770-1740 Ma age for the
regional D2 event (Dahl et al., 1999a, 1999b, 2000). Recent work by Dahl et al.
(2005a,b), suggested a circa 1775 Ma date for the D2 event.
Foliation
The dominant planar feature within the study area is the S2 foliation. The
S2 foliation is also the most pervasive fabric element in the metamorphic rocks of
the Black Hills. S2 is a locally penetrative, north- to northwest-striking, steeply
dipping foliation that is axial-planar to steeply plunging isoclinal F2 folds. This
axial-planar foliation has been designated S2 because an earlier fabric
(presumed S1) is preserved as oblique inclusion trails in regionally
metamorphosed garnets (Dahl et al., 1999a; 2005a,b). However, as noted
above, it is not clear whether this reported S1 fabric is associated with D1,
represents the early stages of D2, or is preserved bedding (S0).
S2 is most easily distinguished in the noses of F2 folds (Figures 12 and
13). S2 is a composite foliation in the limbs of the F2 folds, where it typically
parallels earlier compositional banding or bedding. S2 typically ranges from a
continuous, slaty cleavage in more micaceous units to a spaced, crenulation
cleavage in more quartz-rich units. S2 may also be recognized as an
anastomosing slaty cleavage in some instances. Microscopically, S2 may be
48 A.
B.
Figure 12 - (A) Photographs of F2 fold in mica-quartz schists showing S2 axial planar cleavage. Pencil is parallel to S2 and hammer is parallel to the plunge of the F2 fold. L2 intersection lineations (see text below) is evident above the hammer handle. (B) Overlay showing F2 fold (red), S2 foliation (blue), and L2 lineation (yellow). Hammer handle is parallel to the S2 foliation intersection with folded S0. Pencil is 14 cm long. Hammer is 32.5 cm long. Fault-block H.
49 A.
B.
Figure 13 - (A) Outcrop photographs of F2 fold showing S2 axial planar cleavage in folded mica- quartz schists. F2 fold has nearly vertical plunge and steeply dipping axial surface. Hammer handle is parallel to the S2 foliation direction. Note calc-silicate boudins marking the limbs of a folded layer just to the right of hammer head. Pencil is 14 cm long. Hammer is 32.5 cm long. (B) Overlay indicating compositional layering outlining the fold (red) and trace of S2 foliation. Note the thickened hinge zones and attenuated limbs of the F2 folds and the refraction between different compositional layers. Fault-block H.
50
Figure 14 - Photomicrographs of S2 foliation demonstrating its variable nature. (A) Spaced S2 cleavage in fine-grained graphitic metasiltstone. (B) Overlay of A showing compositional layering (Sc) in red and S2 in blue. (C) Spaced S2 defined by preferred grain orientation and elongated quartz stringers in metapsammites. (D) Overlay of C showing S2 in blue. Note the sheared compositional banding (red arrow). Photomicrographs A-D are standard thin-sections (27 x 46 mm).
A.
B.
51 C.
D.
52 recognized as a preferred orientation of grain boundaries and as the bounding surfaces of microlithons or shear zones at a high angle to bedding or compositional banding (Figure 14).
Variations in S2 and Related Structures Across the Study Area
The orientation of S2 varies among the fault-blocks (Figure 5). Fault- blocks A, B, and C show markedly different structural patterns than fault-blocks D
– I (Figures 15, 16, and 17). Fault-block A is strongly overprinted structurally by the Harney Peak Granite intrusion, and the orientation of the dominant foliation in fault-block A is directly related to the emplacement of the Harney Peak Granite.
It is not clear how many earlier fabrics existed in fault-block A prior to the Harney
Peak Granite intrusion, but all earlier fabrics have been flattened and reoriented by the intrusion. Fault-block B is separated from block A by a cryptic structure, called the Elk Mountain lineament by Ratte and Wayland (1969), in the adjacent
Hill City quadrangle (to the west). I interpret the lineament to be a fault. Rocks in fault-block B are unaffected by the Harney Peak Granite, likely indicating a different deformational history for this block. Fault-block B contains no outcrops of granite and/or pegmatite, and earlier fabrics show no modification by the intrusion. The southern half of fault-block C shows an overall flattening of pre- existing, steeply dipping fabrics by the emplacement of the Harney Peak Granite, while the northern portion of the fault-block is little affected. The southern portion of fault-block C also contains many of the better known pegmatite mines (e.g.
Etta, Hugo, etc), and the overall flattening of S2 may be
53 Equal Area N (Schmidt)
22%
20%
18%
16%
14%
12%
10%
8%
6%
4%
2%
Equal Area N A. N = 93 (Schmidt)
8%
6%
4%
2%
Equal Area N (Schmidt)
B. N = 61
6%
4%
1 2% 3
2
C. N = 104
Figure 15 - Lower hemisphere equal area projections (intensity per 1% area) of contoured poles to dominant foliation and axial-planes of F2 folds (black squares). (A) Fault-block A (S2). (B) Fault-block B (S0/Sc). (C) Fault-block C (S2/Sc). Overlayed numbers 1 and 2 are interpreted to be the limbs of recumbent F2 folds while 3 is interpreted to be Harney Peak deformation. See text for explanation. Note the overall flattening of S2 foliation in (A) and overall shallowly dipping fabric in (B). Fabric in B may not correspond to the regional S2 foliation, although it is the second fabric recognized in fault-block B. This fabric may be S1, and if so, this indicates a different tectonic history for fault-block B. See text for discussion.
54
Equal Area N (Schmidt)
18%
16%
14%
12%
10%
8%
6%
4%
2%
A. N = 287 Equal Area N (Schmidt)
18%
16%
14%
12%
10%
8%
6%
4%
2%
B. N = 131
Figure 16 - Lower hemisphere equal area projections (intensity per 1% area) of contoured poles to S2 foliation and axial-planes of F2 folds (black squares). (A) Fault-blocks D and E. (B) Fault- blocks F and G. Note the slight variation in strike direction between fault-blocks but the pattern remains very similar overall. The pattern in (A) is dominated by a map-scale F2 fold that is evident on the map of the study area (Plate I) in fault-block E. Fault-blocks F and G show a very strong linear fabric. Overall patterns are very dissimilar to fault-blocks A – C (cf. Figure 15).
55 Equal Area N (Schmidt)
14%
12%
10%
8%
6%
4%
2%
A. N = 165 Equal Area N (Schmidt)
20%
18%
16%
14%
12%
10%
8%
6%
4%
2%
B. N = 28
Figure 17 - Lower hemisphere equal area projections (intensity per 1% area) of contoured poles to S2 foliation and axial-planes of F2 folds (black squares). (A) Fault-block H. (B) Fault-block I. Overall pattern of steeply dipping foliation remain consistent between the fault-blocks, while the strike orientation becomes progressively more westward oriented in a west to east transect across the fault-blocks D – I. Overall pattern is again dissimilar to fault-blocks A – C.
due to intrusion of late pegmatites rather than intrusion of the Harney Peak
Granite. This portion of block C dominantly consists of micaceous schist, while
the northern portion is mostly medium- to thick-bedded metagraywacke. Thus,
the discontinuity of the Harney Peak Granite overprint may be related to
differential rheologic susceptibility of the two rock lithologies. Fault-blocks D
56 through I display S2-dominated north-northwest trending, steeply dipping fabrics,
whose strike directions vary only slightly from fault-block to fault-block but rotate
systematically from NNW to NW moving eastward from D-I (Figures 15 to 17).
Folds
F2 folds are the dominant ductile structures in the Mount Rushmore
quadrangle. Outcrop-scale F2 folds in fault-blocks D – I are characterized by tight to isoclinal folds that are vertical to steeply plunging and have axial planes that
are approximately parallel to S2 in the respective fault blocks. F2 folds are
generally northwest trending and have thickened hinges and attenuated limbs
diagnostic of passive flow folds (cf. Ramsay, 1967). Fault-block B contains folds
that otherwise resemble F2 folds, but are tight to isoclinal and are recumbent.
These folds generally lack a distinct S2 foliation or S2 is poorly defined, making
them distinct from typical F2 folds. Figures 12 and 13 illustrate typical outcrop-
scale F2 folds observed in the study area.
Map-scale F2 folding is easily recognized on the published 1:250,000-
scale map of the Black Hills (cf. Figure 7; Dewitt et al., 1986). A particularly good
example is evident in the north-central portion of the Mount Rushmore
quadrangle in the outcrop pattern of banded iron formations (Plate I; fault-block
E). Other examples of F2 folding are not especially well preserved in the map
patterns, most likely due to translation and shearing along the many fault-zones.
In the southwest portion of the quadrangle (fault-block A), the emplacement of
the Harney Peak Granite has obliterated any evidence of F2 folds.
57 Equal Area N (Schmidt)
6%
4%
2%
A. N = 152
Equal Area N (Schmidt)
12%
10%
8%
6%
4%
2%
B. N = 73
Figure 18 - Lower hemisphere equal area projections (intensity per 1% area) of poles to F2 fold axial planes measured during this study. (A) F2 folds in fault-blocks D – I. F2 folds are northwest trending tight to isoclinal folds with steeply plunging hinge-lines. Axial planes are approximately parallel to the dominant S2 foliation. (B) Tight to isoclinal, recumbent folds that otherwise resemble F2 folds in fault-blocks A-C. These folds have shallowly plunging to horizontal hinge- lines and generally have a weak axial-planar fabric. They are similar to the fold shown in Figure 11. One interpretation is that these are early F1 or F2 folds that have been reoriented by the Harney Peak Granite intrusion. Regardless, these folds do not show the strong overprint of D2 and later deformations and have experienced a different deformation history than the folds in fault blocks D – I.
58 Stereographic projections of F2 fold data collected during this study are shown
in Figure 18. Axial plane orientations by fault-block are also shown in Figures 15,
16, and 17. The F2 folds in fault-block B have shallowly plunging to horizontal hinge-lines and generally have a weak axial-planar fabric. They are similar to the fold shown in Figure 11. One interpretation is that these are early F1 or F2 folds
that have been reoriented by the Harney Peak Granite intrusion. Regardless,
these folds do not show the strong overprint of D2 and later deformations and have experienced a different tectonic history.
Lineations
L2 lineations are the most prevalent linear fabric observed in the Mount
Rushmore quadrangle. Fabric diagrams of L2 lineations are shown in Figures 19
and 20. L2’s are generally intersection lineations between bedding/Sc and axial-
planar foliation (S2) that are parallel to the hinge-lines of the steeply plunging F2
folds (Figure 17) and lie within the plane of S2. L2 lineations may also be defined
by elongated minerals such as quartz and biotite in certain lithologies. Elongate
quartz pebbles with long axes oriented in the L2 direction are also recognized in
certain mica-schists. Figure 21 shows an excellent example of L2 defined by elongate quartz rods. This quartz mineral lineation appears to be both an intersection lineation and an elongation lineation. Based on the concentrated distribution of elongation lineations, the eastern part of the study area must have undergone considerable constrictional strain, resulting in LS-tectonites in some domains. However, in other areas the fabric is dominantly an SL-tectonite that
59
Equal Area N (Schmidt)
Equal Area N (Schmidt) A. N = 30
Equal Area N (Schmidt)
B. N = 24
14%
12%
10%
8%
6%
4%
2%
C. N = 320
Figure 19 - Lower hemisphere equal area projections of L2 lineations and F2 hinge-lines measured during this study. (A) L2 data for fault-blocks A (blue circles), B (red squares), and C (black crosses). (B) L2 (black crosses) and F2 hinge-lines (red squares) for fault-block C. (C) Contoured (intensity per 1% area) projections of L2 lineations and F2 hinge-lines for fault-blocks D – I. Note the agreement of L2 and F2 hinge-lines indicating that L2 in fault-blocks D – I lie in the S2 plane. Comparison with figure 18b suggests that the majority of the lineations for fault-block B are parallel to the hinge lines of recumbent folds.
60
Equal Area N Equal Area N (Schmidt) (Schmidt)
18%
14% 16%
12% 14%
10% 12%
8% 10%
6% 8%
4% 6%
2% 4%
2%
A.N = 87 B. N = 72
Equal Area N Equal Area N (Schmidt) (Schmidt)
20%
18%
16%
14%
12%
10%
8%
6%
4%
2%
C.N = 108 D. N = 20
Figure 20 - Lower hemisphere equal area projections (intensity per 1% area) of L2 lineations (contoured) and F2 hinge-lines (black squares). (A) Fault-blocks D and E. (B) Fault-blocks F and G. (C) Fault-block H. (D) Fault-block I (L2 = black crosses; F2 hinge-lines = red squares). Note that orientations of L2 and F2 hinge-lines become more parallel from west to east across the fault- blocks.
61
Figure 21 - Excellent example of L2 lineations defined by elongate quartz minerals in quartz-mica schist. Note the vertical orientation of the L2 lineations on this surface. The orientation of these lineations is identical to the Sc-S2 intersection. Picture from same locality as Figures 12 and 13. Pencil is 14 cm long. See text for explanation.
results in prominent intersection lineations between the S-component of the fabric and compositional banding or bedding. In general, the L2 mineral lineations appear to have developed by elongation parallel to the F2 fold axes.
62 Variation of D2 Deformation Features by Fault Block
Deformational features vary among the fault-blocks of the Mount
Rushmore quadrangle. Fault-blocks A and B show markedly different
deformational features from fault-blocks D – I, and from each other. Fault-block
C shows deformational features that range in orientation from those in fault-block
A to those in fault-blocks D – I. Fault-block A is dominated by the deformational
and metamorphic overprint of the Harney Peak Granite (discussed below). Any
earlier fabrics that may have existed in fault-block A have been transposed
(Figure 15). It is not clear to what extent earlier fabrics may have been present.
Fault-block B contains two fabrics, original bedding and a weak to moderately developed cleavage, which is axial planar to recumbent to reclined isoclinal folds
(e.g. Figure 11). I interpret these folds to be F2 folds that were subsequently
reoriented and flattened by the emplacement of the Harney Peak Granite.
Strong S2 fabrics are not developed in this block and lineations are subhorizontal
and parallel the hinge-lines of the F2 folds (cf. Figure 15b).
Block C shows dominantly S2 fabrics but is locally overprinted by Harney
Peak Granite-related deformation where late pegmatites have intruded the local
country rocks. S2 fabrics in fault-block C are slightly less steeply dipping than S2
in fault-blocks D – I and show correspondingly more scatter (Figures 15 – 20).
The dominant S2 fabric strikes slightly west of north and dips steeply northeast.
A subset of the S2 fabric strikes more westerly and dips variably to the north. F2
folds and L2 lineations are likewise variable within fault block C. In general, F2
axial planes and the S2 fabric show similar scatter (Figure 15c). However, the
63 three concentrations of contoured poles to S2 foliation represent the two limbs of
F2 folds and reorientation of S2 fabrics by Harney Peak related deformation. L2 lineations and F2 fold hinge-lines show tremendous variability in orientation within
fault-block C (Figure 19a, 19c). I attribute this variability of D2 features to the
transitional nature of deformation in fault-block C. For instance, early formed D2 features may have been variably reoriented by late stage pegmatite bodies. In fact, most Harney Peak related deformation in fault-block C is concentrated around localized pegmatite bodies in the southern portion of the fault-block.
Another possibility is that the variation in D2 deformational features within
fault-block C was caused by the intrusion of the Harney Peak Granite. If the
Harney Peak Granite was being forcefully emplaced, and assuming that fault-
block C has had very little strike-slip movement post intrusion, one would expect
to see a north-striking, shallowly dipping reorientation of S2. Any intrusive bodies
postulated underneath the southern half of fault-block C should show similar
reorientation of S2 as shown in fault-block A. It is difficult to envisage how emplacement of an intrusive body immediately west of rocks containing west- striking fabrics (S2a discussed below) could account for the generation of these
fabrics. The two subsets in the S2 fabric could represent the incomplete
transposition of early S2 fabrics, with the more westerly (group 2, Fig. 15c)
striking being the earliest fabric, S2a, and the steeply northeast dipping, more
northwest striking subset (group 1, Fig. 15c) representing late S2b development
and transposition of S2a during oblique convergence, as discussed below.
64 Incomplete transposition of S2a fabrics within fault-block C could have been
accommodated by variation in localized strain partitioning during D2.
Fault-blocks D – I show dominant D2-related deformational features,
although the orientation of these features vary somewhat among the fault-blocks.
Despite the slight variation of S2 among fault-blocks D – I, the overall pattern
remains very similar. The steeply dipping foliation remains consistent between
fault-blocks, while the strike orientation becomes progressively more westward
oriented in a west to east transect across fault-blocks D – I. In all fault-blocks, S2 is generally consistent with F2 axial planes with the exception of fault-block I
(Figures 16 and 17). Likewise, L2 lineations and F2 hinge-lines are consistent in
all fault blocks (Figures 19b, 20). The pattern of fault-block E demonstrates a
map-scale F2 fold that is evident in the quadrangle. From west to east, the fault
blocks show progressively stronger linear fabrics as shown by the stronger
correlation between L2 and F2 hinge-lines.
Faults
Six major faults have been identified in the Mount Rushmore quadrangle
(Figure 5). Multiple smaller fault-zones exist but were not mappable at the scale
of this study, were not extensively traceable, or have little discernable
displacement (Figure 22). The major faults trend north-northwest across the study-area, roughly parallel to the S2 foliation. From southwest to northeast the
faults are: Elk Mountain, Empire Mine, Keystone, Silver City (aka Silver Mtn),
Echo Peak, and Storm Hill.
65 The Echo Peak fault separates fault-block F from fault-block H and
separates quartz-mica schist from staurolite bearing quartz-mica schist. It is
named herein for its exposure at the summit of Echo Peak. The Echo Peak fault
is correlative with an unnamed fault first recognized southeast of the Mt.
Rushmore quadrangle in the Rapid City West quadrangle by Dheeradilok (1974).
The major faults are vertical to subvertical (Figure 23). The Empire Mine,
Keystone, and Silver City faults are also characterized by the presence of
graphite.
Figure 22 - Breccia in quartzite formed during early(?) brittle faulting in the study area. Late brittle faulting (Laramide) also occurred in the study area as the Black Hills were uplifted. Knife is 9 cm long.
Displacement and sense of movement along the majority of the faults are generally indeterminate because of the lack of appropriate marker units and other shear sense indicators. However, where recognized, shear sense indicators (e.g. Figures 24, 25, and 26) give an overall left-lateral sense of
66 A.
B. Figure 23 - (A) Expression of the Keystone Fault in metagraywackes approximately 1.5 km west of the Keystone “Wye” on the north side of US16. Note silicified zone on left of photograph. (B) Same as A with overlay showing extent of fault zone. Note the vertical nature of the fault zone. My wife, Kristin, in the foreground for scale. She is 170 centimeters tall. View to NNW.
67 A.
B.
Figure 24 - (A) Photomicrograph of graphitic rocks from station MR270 showing left-lateral shear as defined by the asymmetry of the porphyroclasts. This sample was taken from the transition zone between the Empire Mine and Keystone faults. (B) Overlay of A outlining some of the more asymmetrical porphyroclasts. Note folded quartz layers in right-center of the photomicrograph. Field of view is approximately 3.5 x 6 mm.
68 A.
B.
Figure 25 - (A) Photomicrograph of graphitic metasiltstone from station MR349 demonstrating left-lateral shear as defined by the asymmetry of the mica-quartz aggregates. Field of view is standard thin-section size (27 x 46 mm). (B) Overlay of A outlining the mica-quartz aggregates.
69 A.
B.
Figure 26 - (A) Photomicrograph of sample MR675, a sheared metapelite collected near the Silver City fault in fault-block H. Note the intense shearing and included garnets (green arrows) with clear rims indicating two-stage growth or change in growth conditions. Late staurolite porphyroblasts (red arrow) overgrow the dominant foliation in the center of the photograph. (B) Photomicrograph of MR733 in the Silver City fault zone. Note the complete transposition of earlier fabrics in B. The late skeletal garnet (red arrow) in B is overgrowing the transposed foliation and the thin-section is graphite rich. Quartz-rich layers above the garnet are shear- folded (yellow arrow). Field of view is 27 x 46 mm for both A and B.
70 movement parallel to the dominant S2 foliation. Moreover, the overall map
pattern along the eastern edge of the Precambrian Black Hills indicates an
overall pattern of strike-slip faulting that is especially evident in the metabasites and related rocks (cf., Figure 7).
Evidence presented in this study indicates that major faulting was initiated
prior to the emplacement of the Harney Peak Granite during the later stages of
D2 deformation. Based on cross-cutting relationships, faults recognized in the
eastern part of the study area must be younger than similar structures west of the
study area, which have been deformed by the Harney Peak Granite (e.g. Hill City
fault, Grand Junction fault, Burnt Fork fault) (Figure 27). Overgrowth of fault
zones by late M2 porphyroblasts (Figure 26) also indicate that major faults
predate the emplacement of the Harney Peak Granite.
Recognition of a Major Structural Discontinuity
A major structural break exists along the north-eastern boundary of fault-
block C (Keystone fault; Figure 27) across most of the Mount Rushmore
quadrangle. In the south central portion of the quadrangle, the Keystone fault is
bifurcated (Plate I). This bifurcated zone is an area of intense deformation.
Multiple rock types are juxtaposed along the 3 km length of this zone. The rocks
are heavily altered and are extremely graphitic in much of the zone. Rock types
include quartz-mica schists, graphitic schists, banded-iron formations, and
quartzite. Along the west-central portion of the study area, the structural
discontinuity is marked by the Empire Mine fault as displacement is seemingly
71 transferred from the Keystone fault to the Empire Mine fault. The Empire Mine fault shows progressively more deformation to the northwest as the Keystone fault shows correspondingly less deformation. Along the western boundary of the quadrangle, the Empire Mine fault shows areas of intense deformation and alteration similar to that along the more southern portions of the Keystone fault described above. The Empire Mine fault continues northwest into the Hill City and Silver City quadrangles where it truncates all earlier structures (Figure 27).
The overall sense of movement along the structural discontinuity is sinistral (cf.,
Figure 24). The direction of movement is constrained by C- and S-type shear band cleavages, offset of lithologic marker units (Plate I), the asymmetry of veins, and sheared porphyroclasts adjacent to the fault zone (Hill et al., 2004). The overall pattern of mapped faults in the eastern part of the Precambrian Black Hills also supports an overall strike-slip motion (cf. Figure 7). This pattern is especially evident in the offset of metabasalts, amphibolites, and metagabbros mapped in the northeast Black Hills. Nabelek et al. (in press) have also demonstrated that the regional M1 biotite metamorphic isograd is more or less
coincident with the trace of the Silver City fault (Figure 6). This suggests that the
Silver City fault offsets regional M1 metamorphic isograds, again indicating significant displacement along the Silver City fault.
It is clear that the faulting along the eastern portion of the Black Hills must be younger than that in the south-central portion west of the Empire Mine fault based on cross-cutting relationships. The Empire Mine-Keystone faults truncate all other structures indicating that faults to the east are younger. Folding of the
72
Figure 27 - Outline map of the Precambrian Black Hills uplift showing locations of major faults mentioned in this study. Note that the Empire Mine-Keystone faults truncate all other structures indicating that faults to the east are younger. Also note the folding of the Grand Junction and Hill City faults by the Harney Peak Granite suggesting that the intrusion post-dates faulting to the west. Compare with Figure 7.
73 Grand Junction and Hill City faults by the Harney Peak Granite suggests that the
intrusion post-dates faulting to the west of the intrusion. (Figure 27; cf. Dewitt et. al, 1986). What remains unclear is whether any of the faults are related to the D1
event. Folded faults that follow lithologic contacts would suggest a D1
relationship. However, the intrusion of the Harney Peak Granite as a late syn- to
post- deformational pluton has also resulted in folding of faults in the central
Black Hills. Nabelek et al. (in press) favor late syn-deformational intrusion of the
Harney Peak Granite and uplift along the Hill City and Empire Mine-Keystone
fault systems based on precipitation of graphite along the fault zones, secondary fluid inclusion data, and overall structural discontinuity along the bounding faults of the Harney Peak block.
My observations indicate that most major faults in the Black Hills are related to the D2 deformational event and that the D2 event encompasses both
the “Black Hills” and “Dakotan” orogenies of Chamberlain et al. (2002).
Approximately 50 million years separate the two “orogenies” based on published
age data (e.g., Dahl et al., 1999 a,b; Chamberlain et al., 2002; Dahl et al.,
2005a,b). As such, they can be thought of as a single, protracted event, which I have divided into D2a and D2b phases. D2a resulted in the north-directed
thrusting and early folding and faulting. D2b structures formed as collision
became progressively transpressive. With respect to this nomenclature, faults
like the Grand Junction fault and the Hill City fault are related to D2a while the
Empire Mine-Keystone fault and all faults east of the structural break are related
to D2b. A tectonic model using this nomenclature and its relationship to
74 structures found in the Mt. Rushmore quadrangle and in the eastern Black Hills is
discussed in greater detail later in this study.
D3 Deformation
D3 deformation was not recognized in the study area, but is recognizable
from outcrop patterns north of the Mt. Rushmore quadrangle in the Silver City,
Deerfield, and Pactola Reservoir quadrangles (Redden et al., 1990; Redden et
al., 2005; Redden, unpublished mapping). Noble et al. (1949) recognized F3 folds in the Homestake mine area and referred to them as “cross-folds”. F3
folding affects much of the central Black Hills (Redden et al., 1990). F3 folds are
doubly plunging folds that refold F2 folds (Redden et al., 1990). Redden (Redden
et al., 1990; Redden, 2001, pers. comm.) describes these folds as “cross-folds”
or “porpoises.” These folds are not common in the southern Black Hills, but may
be recognized on more detailed maps north of the Mount Rushmore quadrangle.
Dahl et al. (1998) suggested that renewed convergence along the southern margin of the Wyoming craton may have produced these folds. They are generally restricted to single fault-blocks and are oriented somewhat obliquely to the bounding faults.
An alternate possibility is that these F3 “cross-folds” represent en echelon
arrays of folds formed as subsidiary structures during strike-slip movement of the
bounding faults. The acute angle between the F3 cross-folds and the bounding
faults would indicate the direction of movement for the adjacent fault-blocks,
which based on unpublished mapping of Redden (pers. comm.) is consistent with
overall sinistral translation of the fault-bounded blocks. F3 folds could also
75 represent shear or possibly drag folding of earlier F2 folds during translation of
the containing fault-blocks. Fabric data collected in fault-block B may also
support the idea of shear folding during translation of the fault blocks that
resulted in F3 cross-folds. Measured S0 and Sc attitudes within fault-block B
record markedly different structural trends based on their location with respect to
the trace of the Empire Mine fault (Figure 28). Those attitudes taken closest to
the fault have a mean principal orientation of 285/58 and a beta axis of 05/102,
while attitudes taken away from the fault have a mean principal orientation of
034/65 and a beta axis of 01/214. This difference may reflect a counter-
clockwise rotation (i.e. drag folding) of material closest to the fault zone during
left-lateral strike-slip movement along the Empire Mine fault.
D4 Deformation
D4 deformation is primarily restricted to fault-block A, which contains the
Harney Peak Granite, and the immediately adjacent southern half of the east- adjacent fault-block C (Figures 5 and 6). D4 is associated with the intrusion of
the Harney Peak Granite. Multiple pegmatites of varying size (one meter or less)
intrude the country rocks over the southern third of the Mt. Rushmore
quadrangle, but only locally deform the country rocks. The main body of the
Harney Peak Granite and its satellite plutons occupy an area of roughly 100 km2 in the south-central Black Hills. The mechanisms of emplacement of the main
Harney Peak Granite have been inferred by analogy from the study of one of its smaller satellites (~1 km2), the Calamity Peak pluton (Duke et. al, 1988). The
76 model suggests that the Harney Peak Granite consists of multiple pulses of
genetically related magma that were initially discordant sills fed by concordant
feeder dikes. Each successive pulse was injected forcefully along the upper
contact between pre-existing granite layers and the country rock. Continued
growth by forceful injection resulted in ballooning of the granite and upward and
outward growth of the granite and produced an overall dome-shaped structure of the Harney Peak Granite and associated D4 fabrics.
Foliation
The forceful emplacement of the Harney Peak Granite resulted in a
roughly concentric fabric (S4) adjacent to the pluton, strong crenulation
cleavages, and transposition of all earlier fabrics (Figure 29; c.f., Figure 15a).
Note the overall similarity of S4 to reoriented S2 fabrics. S4 fabrics generally
show variably developed D4 structures such as a spaced or crenulation cleavage
that overprint D2 deformational features as a result of forceful emplacement of
the Harney Peak Granite. The two maxima in this diagram likely represent the
change of S4 orientation along the curved periphery of the Harney Peak Granite.
The pattern looks very similar to that of fault-block B (Figure 15b) indicating that
the Harney Peak Granite is flattening earlier F2 folds in fault-block A.
The southern portion of fault-block C immediately adjacent to the Harney
Peak block also shows some deformation that is likely associated with D4. As discussed previously, the larger economic pegmatites that occur in this block and the occurrence of schists may account for the local deformation. Locally, smaller
77
Figure 28 - Lower hemisphere equal area projections of poles to S0/Sc foliation in fault-block B. (A) S0/Sc measured away from the trace of the Empire Mine fault. (B) S0/Sc measured in close proximity to the trace of the Empire Mine fault. Note that a counter-clockwise rotation of material closest to the fault could explain the difference and indicate possible initiation of drag folding not recognized in the field. See text for explanation. pegmatites deform the rocks into which they are intruded but not on a large scale. Elsewhere in fault-block C and all other fault-blocks, rocks are unaffected by the emplacement of the Harney Peak Granite.
78 Equal Area N (Schmidt)
16%
14%
12%
10%
8%
6%
4%
2%
Lower Hem. N = 79
Figure 29 - Lower hemisphere equal area projection (intensity per 1% area) of poles to S4 foliation measured in fault-block A.
Faulting
No evidence for distinct faults that can be directly related to D4 was
observed in the Mount Rushmore quadrangle. However, Nabelek et al. (in press) suggest that the Harney Peak block was uplifted along the Hill City and Empire
Mine-Keystone faults based on the presence of late andalusite in the metamorphic aureole and the sharp difference in the dominant foliation. They suggest that the Harney Peak block was initially intruded by magma at a greater depth than its current relative position and was uplifted to its present structural level. Fluid inclusion data from pegmatites suggest late, low pressure deformation (Sirbescu and Nabelek, 2003), suggesting post-intrusive movement along the faults. At the very least, movement along the major fault systems
79 continued after the intrusion of the Harney Peak Granite (Hill et al., 2004;
Nabelek et al., in press).
D5 Deformation
A post intrusion D5 event, which produced a northeast-trending fabric
across the central Black Hills ca. 1535 Ma, was first suggested by Redden et al.
(1990), based on the slope of a bulk sample Rb-Sr isochron plot for whole-rock
samples of the Montana Mine formation and metagraywackes near the Three
Forks area, 8 km northeast of Hill City. Redden described locally recognizable
NE-trending D5 related folds with a spaced axial-planar cleavage. Distlehorst
(1999) reported a weakly developed, northeast-southwest striking, spaced biotite cleavage in the Rockerville area, west of the study area.
Foliation
In the study area, the D5 event is recognized as spaced cleavage (S5) or, less commonly, as a late crenulation cleavage (Figures 30 and 31) that cross- cuts the dominant S2 foliation at a high angle. In the southeastern portion of the
study area, S5 is a crenulation cleavage that progresses to a spaced cleavage to
the east. This crenulation was first recognized in the center of the quadrangle
east of the Silver City fault and may be seen in the road cuts north of US-16 beginning at the intersection of US-16 with the mapped trace of the Silver City
fault (Plate I). This crenulation cleavage (Figures 31 and 32) progresses to a
slaty cleavage eastward and is recognized as such by the eastern boundary of
80 the quadrangle. Because this fabric intensifies away from the Harney Peak
Granite, it is unlikely that it is in any way related to the intrusion. Microscopically,
D5 may be recognized as either a spaced cleavage, defined by the orientation of
retrograde biotites or chlorite porphyroblasts, or less commonly as a crenulation
cleavage (Figures 33, 34).
Folds
One probable small-scale D5 related fold was identified in the quadrangle
(Figure 35). This fold had an axial plane orientation of 020/81 with a plunge of
70/197. The measured S5 foliation was not axial-planar to this fold with an
orientation of 212/66, although it appeared to be in the outcrop. The attitude of
this fold is consistent with described F5 generation folding (e.g., Redden et al.,
1990). No other likely F5 folds were recognized in the Mount Rushmore
quadrangle even though the S5 fabric was evident over much of the southeastern
portion of the study area.
Summary of Deformational Events
Three major deformational events are recognized in the study area. Table
3 summarizes events recognized in this study and compares them to previous
studies. None of the fault-blocks in this study area exhibit every deformational
fabric. Fault-blocks A, B, and C show very distinct structural histories. Fault- blocks D – I are similar with respect to each other but are markedly distinct from
fault-blocks A, B, and C. A major structural break exists along the
81 A.
B.
Figure 30 - Field photographs showing the well developed S5 spaced cleavage and its relationship to the dominant S2 fabric. (A) Photograph of S5 recognized in the limb of a F2 fold. S5 closely spaced and is nearly vertical and S2 is parallel to the pencil and trends left to right. Pencil is 14 cm long. (B) Photograph of wider spaced S5 on the S2 surface. S5 trends left to right in the photograph and S2 is nearly vertical. Knife is 9 cm long.
82 A.
B.
Figure 31 - (A) Field photograph of S5 crenulation cleavage developed east of the Silver City fault. Knife body is parallel to S2 (nearly vertical) and knife blade is parallel to S5. S5 is plunging steeply in the viewing direction. Body of knife is 9 cm long. Blade of knife is 7.5 cm long. (B) Spaced S5 cleavage (trends bottom right to top left) cross-cuts S2 foliation (nearly vertical and parallel to pencil). Pencil is 14 cm long.
83 A.
B.
Figure 32 - (A) Photomicrographs of S5 as defined by bounding surfaces of microlithons. (B) Overlay of E showing S5 in blue. Field of view for photomicrographs A and B is approximately 1 x 1.5 mm. Thin section of sample MR945, fault-block F.
84 A.
B.
Figure 33 - Photomicrographs of S5 foliation. (A) Sample MR932 showing distinct S5 spaced cleavage (nearly vertical) defined by biotite selvages. S2 is nearly horizontal. Note the folded and transposed quartz stringers that define the S2 fabric. (B) Sample MR740 showing less distinct S5 spaced cleavage and aligned chlorite porphyroblasts. S2 is subhorizontal and defined by the preferred orientation of small, earlier biotites. Field of view is 27 x46 mm.
85
Figure 34 - Photomicrograph of sample MR915 showing well developed S5 crenulation cleavage trending subvertically right to left. S2 is subhorizontal. Field of view is 27 x 46 mm.
Figure 35 - Photograph of possible F5 fold showing warping of S2 foliation. A weak spaced foliation (S5) is parallel to the pencil trending bottom right to top left. S5 appears axial planar but dips in opposite direction. Pencil is 14 cm in length.
86 north-eastern boundary of fault-block C (Keystone fault) across most of the study
area. Along the west-central portion of the study area, the structural discontinuity
is marked by the Empire Mine fault as displacement is seemingly transferred
from the Keystone fault to the Empire Mine fault (i.e., possible left stepover). The
Empire Mine fault continues northwest into the Hill City and Silver City
quadrangles where it truncates all earlier structures (cf. Dewitt et al., 1986;
Figure 7). My observations are consistent with the assertions of Hill et al. (2004)
and Nabelek et al., (in press) that fault-blocks in the study area and, by
extrapolation, east of the structural break have differing structural and P-T
histories.
TECTONIC OVERVIEW AND OBLIQUE CONVERGENCE MODEL
The history of the assembly of southern Laurentia is an epic involving
multiple Archean and Paleoproterozoic landmasses undergoing simultaneous
convergence along multiple margins. Although the current state of knowledge is
adequate to decipher the relative timing of these events, much debate continues.
Current and future research will undoubtedly refine both the absolute timing and tectonic history of southern Laurentia. Much work remains to be done and the tectonic synthesis presented herein is knowingly inadequate. It is beyond the scope of this paper to discuss in detail the tectonic history of southern Laurentia.
87 Folds Fabrics Faults
Style Orientation Remarks
*not observed in
D1 this study
*possible Early, weak axial-planar Grand Junction, Hill City, correlation with D Reclined to as observed in fault- 1 Tight to Burnt Fork, and other structures of recumbent; NE block B; possibly related D2a isoclinal faults west of structural Redden et.al (1990) trending to observed break and Dahl et. al compositional banding (1998a)
North to northwest Empire Mine, Elk Mtn, Dominant S ; steeply trending, 2 Keystone, Silver City, *S of Redden et. al Isoclinal dipping, axial planar 1 D2b upright to Storm Hill; reactivation of (1990) fabric steeply Hill City(?) plunging
*Not observed in Possible drag this study; F3 of folding along Redden et. al
D3 Empire Mine (1990) and "cross- fault; folds" of Noble et.al (1949)
* Uplift of entire Harney Peak block Local schistosity and Uplift along Empire Mine (Nabelek et. al, in crenulation cleavage and Hill City faults (?); press). Not an recumbent; associated with granite Elk Mountain fault; actual "doming" Tight to reorientation of emplacement; Restricted Golden Slipper and event caused by D4 isoclinal D2a folds to fault-block A and Rabbit Gulch brecciated intrusion of Harney southern, adjacent zones (Ratte & Wayland, Peak Granite. portion of fault-block C 1969) Doming and S3 of Redden et. al (1990).
Northeast trending crenulation and spaced *S of Redden et. al cleavage in the southern 4 D5 (1990) portions of fault-blocks E, F, and H
Table 3 - Revised summary of deformational events and related structures based on this study. Compare with Table 2.
My intention here is to provide sufficient knowledge to familiarize readers with the tectonic framework of southern Laurentia, and thus to be able to discuss a tectonic model for the study area and its broader implications.
88 Laurentian Assembly: Overview
J. Tuzo Wilson (1962) first advanced the idea of “mobilism” of Proterozoic
orogenic belts exposed in the Canadian shield, providing the impetus to apply
new techniques and concepts to the study of Precambrian rocks. Over the past
several decades, much has been learned about the assembly and evolution of
Laurentia. An excellent synthesis of geologic, geophysical, and geochemical
data by Hoffman (1988, 1989) outlines the framework regarding Laurentian
assembly. Figure 1 illustrates the generalized tectonic elements involved in the
assembly. The assembly of northern Laurentia, where exposed in the Canadian
shield, is relatively well understood as compared to that of southern Laurentia
where parts of the shield are masked by Phanerozic cover.
In the southern portion of Laurentia, the 1.9- 1.83 Ga Penokean orogeny
occurred along the southern boundary of the Superior Province and was the
result of the accretion of one or more island-arcs formed over a south dipping
subduction zone (e.g., Peterman et al., 1985; Sims and Peterman, 1986;
Hoffman, 1988; Sims et al., 1989, 1993; Van Wyck and Johnson, 1997; Holm et
al., 1998). The northern margin of the Superior Province records the 1.9-1.83 Ga
Trans-Hudson orogeny that resulted from the collision of the Hearne-Rae
Province with the Superior Province. Tectonic polarity for the orogen is interpreted on the basis of the internal zones of the orogen, with the Superior
Province as the foreland and the Hearne Province as the hinterland. The term
Trans-Hudson orogeny has also been used to refer to the 1780 – 1715 Ma east-
west directed collision between the Wyoming and Superior provinces (e.g.,
89 Redden et al., 1990; Dahl et al., 1998, 1999 a, b, 2005a,b; Chamberlain et al.,
2003), however, as discussed below, this designation may not be appropriate.
The collision between the Wyoming Province and the Hearne-Rae Province to
the north has historically been inferred to be pre-Trans-Hudson, however this
interpretation is also equivocal and the discussion below reexamines this inference.
Incongruities with Laurentian Assembly
Although the basic framework is in place to recount the assembly of
Laurentia, many problems and controversies still exist (Lewry et al., 1990;
Machado, 1990; Boone and Hynes, 1990). The building of Laurentia involved at
least six Archean microcontinents: Slave, Rae, Hearne, Wyoming, Superior, and
Nain, which were sutured along multiple Proterozoic orogenic belts (Figure 1;
Hoffman, 1988). Even in the Canadian shield where the Proterozoic belts are
best exposed, complications arising from lack of good age constraints and
imprecise structural relationships yield multiple viable interpretations of the same data (e.g. Hoffman 1990, Lewry and Collerson, 1990).
The complications in elucidating the assembly of Laurentia are most
pronounced along the boundaries between the southern-most Archean
components (Wyoming, Superior, Hearne provinces, and Medicine Hat Block),
where the boundary zones are obscured by Phanerozoic continental platform
cover sequences (Figure 2). Limited exposure of Archean age rocks requires
that boundary zones between Archean age provinces be inferred primarily from
90 geophysical data. However, geochemical studies of the Archean age rocks of
Laurentia (e.g., Wooden and Mueller, 1988; Mueller et al., 1988; Frost and
Chamberlain, 1998; Frost, 2000) demonstrate that these early Archean cratons
are distinct in their tectonic evolution and were likely widely distributed prior to
terminal collision. Hence, the inferred geophysical boundaries coincide with
convergent margins along multiple boundaries.
The most significant collisonal event of the “Pan-American System” (Lewry
and Collerson, 1990) is the Trans-Hudson orogeny, which is the terminal
collisonal event between northern and southern Archean elements of Laurentia.
The Trans-Hudson orogen, as popularly defined, extends from central Greenland
across Hudson Bay and from there west-southwest through the exposed
Canadian shield to its type locality in northern Manitoba and Saskatchewan and is interpreted to continue southward under Phanerozoic cover along the boundary between the Wyoming and Superior provinces (Lewery and Collerson,
1990; Hoffman, 1990).
Problems with “Trans-Hudson”
Hoffman (1990) stated that the name “Trans-Hudson” orogen (Hoffman,
1981) was intended to identify the Early Proterozoic collision zone between the
Superior Province and the Archean domains of the northwest Churchill Province
(i.e., Hearne-Rae provinces, the Nain Province, and intervening juvenile
terranes) and that the name was an afterthought that lacked precise definition.
The term has been applied to the deformed boundary zone between the
91 Wyoming and Superior provinces (Dakota segment) and is engrained in the
literature of the 1990’s (e.g. Redden et al., 1990; Klasner and King, 1990; Nelson
et al., 1993; Baird et al., 1996). Multiple lines of evidence suggest that the
application of the term “Trans-Hudson” to the Wyoming-Superior collisional zone
is inappropriate. The Trans-Hudson orogeny, as popularly defined, is largely
buried over the ~2800 km trace of the orogen. The lack of exposure over the
trace of the orogen necessitates dependence on geophysical methods to define the orogen and interpret the structural complexity. Other enigmatic structures, such as the Great Falls Tectonic Zone (GFTZ) and the Vulcan Structure (cf.,
Mueller et al., 2002, 2005) (Figure 2), have traditionally been interpreted to be separate from the Trans-Hudson. Recent work in the GFTZ (e.g. Mueller et al.,
2002; 2005) suggests that the traditional interpretation is equivocal. Finally, the age and character of the Trans-Hudson in Canada as compared to the Dakota segment is quite different.
Exposure
Only in the central Canadian shield is the Trans-Hudson well exposed
over a broad corridor with strike-length of approximately 500 km, although
sizeable areas have been mapped near the eastern shore of Hudson Bay
(Thomas, 2001). North of the central Canadian shield, the Trans-Hudson
crosses underneath the Paleozoic Hudson Bay basin and then presumably
reappears in Greenland and Baffin Island (Mukhopadhyay and Gibb, 1981;
Hoffman, 1990; Lewry and Collerson, 1990). However, Hoffman (1987, 1988,
92 1990) argued for the exclusion of the Greenland and Baffin Island segments with
respect to the Trans-Hudson, based primarily on the location of possible sutures
and the degree of continuity of juvenile terranes. South of the central Canadian shield, the Dakota segment of the “Trans-Hudson” is buried under Phanerozoic and younger rocks and is only exposed in the ~40 km x 80 km Black Hills uplift
(Figures 2 and 3). Rocks exposed in the Dakota segment are also considerably younger than Hudsonian (cf. Redden et al., 1990; Dahl et al., 1998; McCombs et al., 2003; Dahl et al., 2003; Dahl et al., 2005a,b). Precambrian rocks of
Hudsonian age are, however, exposed within the GFTZ in the Little Belt
Mountains, Montana (cf. Mueller et al., 2002, 2005 and references therein).
Geophysical Interpretation
Extrapolation of lithologic domains in the Trans-Hudson to the Dakota
segment was initially developed largely based on a very short seismic reflection
profile in southern Saskatchewan adjacent to the exposed part of the orogen
(Green et al., 1980, 1985; Klasner and King, 1990; Nelson et al., 1993). Green
et al. (1985) extrapolated terranes mapped within the exposed portion of the
Trans-Hudson orogen in Canada southward along the deeply buried collisional
zone between the Wyoming and Superior provinces based on geophysical
interpretation of regional aeromagnetic, gravity, and seismic reflection and
refraction data. Using aeromagnetic and gravity data along with limited drill core
data, Klasner and King (1986, 1990) further subdivided the buried collisional
zone into five lithotectonic terranes that corresponded to the Lewry et al. (1985)
93 subdivision of the exposed Trans-Hudson in Canada. Klasner and King (1990)
also advanced a crustal model for the Dakota segment (Figure 36) based on the
idealized orogenic model of Hatcher and Williams (1986). Nelson et al. (1993)
presented a summary report of new seismic data collected by the Consortium for
Continental Reflection Profiling (COCORP) approximately 700 km south of the
exposed Trans-Hudson in Canada. Baird et al. (1996) presented a more detailed
analysis of the COCORP data.
The Dakota segment has been ingrained in the literature as “Trans-
Hudson” based primarily upon the interpretations of Camfield and Gough (1977)
and Klasner and King (1986, 1990; Figure 36). However, based on the
COCORP data collected in 1990, Nelson et al. (1993) and Baird et al. (1996)
demonstrated that the earlier extrapolations of lithotectonic domains based on
correlation with exposed parts of the Trans-Hudson are not appropriate (Figure
38). New Lithoprobe (Canada’s national geoscience project) data immediately
adjacent to the exposed Trans-Hudson also rebuts the model of Klasner and
King (1990) (Nelson et al., 1993; Lucas et al., 1993; White et al., 1994). Both the
COCORP data and the Lithoprobe data record structural relations opposite to what was predicted by the existing interpretations of the evolution of the Dakota segment (e.g., Lewry et al., 1990; Klasner and King, 1990).
The new COCORP data (Figure 37) also demonstrates a widely different
MOHO signature for the exposed Trans-Hudson versus the Dakota segment
(Nelson et al., 1993; Baird et al., 1995; Baird et al., 1996). The exposed Trans-
94 Hudson shows a sharply defined subhorizontal MOHO. In contrast, the Dakota
segment shows no evidence of this feature. Dipping reflectors in the Dakota
segment are reported to die away at two-way travel-times appropriate for the
upper mantle. This dissimilarity between the Trans-Hudson proper and the
Dakota segment is significant as sharply defined subhorizontal MOHO’s are
indicative of collapsed (or collapsing) orogens (Nelson, 1991).
Possibly the least well understood geophysical feature associated with
Laurentian assembly is the North American Central Plains electrical conductivity
anomaly (NACP; Figure 38). Extensive studies have shown that the NACP
extends from northern Canada into the Dakotas (e.g., Alabi et al., 1975; Camfield
and Gough, 1975; Jones et al., 1993, 2005; Garcia and Jones, 2005). Camfield
and Gough (1977) suggested that the NACP marks the Proterozoic collisional zone from the southern Rockies to northern Canada. In the Dakotas, the NACP was interpreted to delineate the eastern boundary of the Wyoming Province
(Dutch, 1983; Klasner and King, 1986). Jones and Savage (1986) relocated the
NACP boundary eastward, where it coincides with the eastern boundary of the later discovered Dakota block (Nelson, 1993; Jones et al.,1993; Baird et al.,
1996). The cause and tectonic significance of the NACP is still under investigation (e.g., Jones et al., 2005; Garcia and Jones, 2005; Ferguson et al.,
2005; Evans et al., 2003, 2005). Most workers (e.g., Jones et al., 1997, 2005;
Garcia and Jones, 2005; Ferguson et al., 2005; Evans et al., 2003, 2005)
95
Figure 36 - Tectonic crustal model modified from Klasner and King (Figure 5d, 1990). Purple = Wyoming Province; Gray = Superior craton; Red = Harney Peak type anatectic granite. Light greens= Thrust belts; Medium green = ophiolites; Dark green = Mafic Intrusives; Light tan = Accreted oceanic sediments, volcanics, and continental blocks; Yellow = Archean sedimentary basins.
Figure 37 - COCORP seismic transect across the southern Trans-Hudson orogen beneath the Williston Basin. Upper figure is migrated line drawing from data. Lower figure is interpretation of Baird et al. (1996). Red line is the MOHO. Figure modified from Baird et al. (Figure 2, 1996). Purple = Wyoming Province; Dark Brown = Dakota block; Light tan = marginal basin rocks; Taupe = reworked Archean marginal rocks; Grey = Superior craton.
96
Figure 38 - Location of the NACP conductivity anomaly in Central North America. From Garcia and Jones, 2005.
associate the NACP with sulfide mineralization, not graphite mineralization
(Camfield and Gough, 1977). Geologically, the NACP has been interpreted to be
related to “carbonated and metamorphosed slate rocks” deposited in foreland
basins (Boerner et al., 1996) or sulfide deposits in fold-hinges of volcanic rocks
(Jones et al., 1997, 2005). A characteristic and enigmatic feature associated
with magnetotelluric investigations of the NACP is that it only affects currents
parallel to strike and has little observable effect perpendicular to strike (Jones,
1993). Moreover, magnetotelluric studies of ancient subduction-collision zones
show that some have associated conductivity anomalies while others do not
(Jones et al., 1993). Therefore its tectonic significance in southern Laurentia is
equivocal.
97
Age and Character of “Trans-Hudson” Rocks
The southern continuation of the Trans-Hudson orogen is buried
underneath Phanerozoic cover sequences (Figure 2). The proximity of the Black
Hills uplift to the exposed Trans-Hudson and the trace of the NACP anomaly
(Camfield and Gough, 1977) resulted in the early accepted continuation of the
Trans- Hudson through the Black Hills. Multiple investigations (e.g. Dahl et al.,
1999 a, b; Nabelek et al., 2001; Chamberlain et al., 2002) have since recognized
the incongruity of dates associated with terminal collision of the exposed Trans-
Hudson (~1900-1830 Ma; Sims and Peterman, 1986; Bickford et al., 1990) with
the inferred southern continuation (~1760-1715 Ma; Dahl and Frei, 1998) in the
Black Hills. The Trans-Hudson is broadly coeval with the Penokean, Wopmay, and Ketilidian orogens (Hoffman, 1988), thus implying nearly simultaneous orogeny along multiple margins. Continued tectonic convergence resulting in the later “Black Hills” orogeny however, does not necessarily connote equivalence.
The recognition of the intervening Dakota block (Baird et al., 1996) undermined the detailed correlation of Klasner and King (1986, 1990) between the Dakota segment and the subprovince delineation of the Trans-Hudson orogen in northern Canada. Moreover, the Great Falls Tectonic Zone, generally considered to be the boundary between the Archean Medicine Hat block (Hearne province) and Wyoming province, yields ages more closely related to the Trans-
Hudson proper (1867 +/- 6 Ma 07Pb/206Pb age; Mueller et al., 2002). Mueller et
al. (2002) interpreted the Great Falls Tectonic Zone to represent a convergent
98 margin that developed during the closure of an ocean basin along the northwestern margin of the Wyoming craton at ~1.9 Ga. This strongly suggests a more definite temporal connection with the Trans-Hudson orogen than that of the Dakota segment.
The Trans-Hudson as exposed in Canada also contains a collage of mainly juvenile, arc-related Early Proterozoic terranes (Baird et al., 1996). These arcs, associated inter-arc sedimentary basins, and older crustal elements were intensely reworked during the Trans-Hudson orogeny (Baird et al., 1996).
Geochronological data reveal no consistent age differences between these domains (Lewry et al., 1994). The existence of extensive intervening juvenile terranes in the Dakota segment remains equivocal (e.g. Baird et al., 1996). The lack of juvenile terranes, a primary argument of Hoffman (1988, 1990) to exclude
Greenland and Baffin Island segments from the Trans Hudson, taken with the age disparity suggests a different progression for the Dakota segment and Trans-
Hudson orogen.
Data for mafic to felsic igneous and metamorphic rocks of the Great Falls
Tectonic Zone demonstrate that they were derived from underplated mantle sources and not from older crust (Mueller et al., 2002), again suggesting a stronger connection to the Trans-Hudson orogen than that of the Dakota segment. Mueller et al. (2005) state that the magmatism and metamorphism in the Great Falls Tectonic Zone (~1.86 Ga; Dahl et al., 2000; Mueller et al., 2002) is temporally equivalent to the Trans-Hudson Orogeny. However, Boerner et al.,
(1998) cited the lack of a collinear magmatic arc and the failure of
99 electromagnetic studies to detect a plate-edge foreland basin and interpreted the
Great Falls Tectonic Zone as a reactivated intracontinental suture zone rather than a Proterozoic age suture between the Wyoming and Hearne provinces.
Wyoming Province Orogenies
Multiple workers (e.g., Dahl and Frei, 1998; Dahl et al., 1999 a, b; Nabelek et al., 2002; Chamberlain et al., 2002) have recognized the diachroneity of terminal collisions for the exposed Trans-Hudson orogeny and southern Dakota segment as discussed above. Dahl and Frei (1998) revived the term “Black Hills orogeny” (Goldich et al., 1966) to distinguish the southern Dakota segment from the Trans-Hudson. Chamberlain et al., (2002) suggested two separate events:
(1) the 1.78 Ga Black Hills orogeny and (2) the 1.72 Ga Dakotan Orogeny. The older, ca. 1.78 Black Hills orogeny is coeval with other orogenic events along the southern (Cheyenne Belt; Chamberlain, 1998) and northwestern margins of the
Wyoming Province (Great Falls Tectonic Zone; Mueller et al., 2005).
This study suggests a two-phase, protracted single event along the eastern margin of the Wyoming Province that encompasses both the Black Hills and Dakotan orogenies (cf. Tables 1 and 2). Shackleton and Ries (1984) state that post-collisional movements are responsible for most observed structures in orogenic belts. Although there is much validity to that statement, it is undoubtedly an oversimplification. Many collisional orogens do exhibit multiple distinct deformational phases caused by a single collisional event (e.g.
100 Laubscher and Bernoulli, 1982; Duncan, 1984; Ellis 1986; Vannay and Steck,
1995; Pignotta and Benn, 1999). The Black Hills are a prime example of this.
Transpressional Model for the Study Area
Data from this study are difficult to interpret with respect to the traditional
“Trans-Hudson” model for the Black Hills. Traditionally, the Proterozoic core of the Black Hills has been described as a series of small epicratonic rift basins
(Gosselin et al., 1988). However, this model (c.f. Redden et al., 1990) has been insufficient to explain the complexities of the exposed Precambrian terrane. New and refined age data (e.g., Dahl et al., 2005a,b; McCombs et al., 2003; Dahl et al., 2003; Chamberlain et al., 2002; Dahl et al., 1998) demonstrate that the rocks of the Paleoproterozoic Black Hills underwent a distinctly younger collisional event that is not related to the Trans-Hudson orogen.
Structural patterns observed in the study area and in the east-central and northeastern Proterozoic Black Hills are best explained by an oblique tectonic collision. Orogens must involve some component of oblique convergence as a result of basic plate kinematics and relative plate motion on a spherical surface.
Oblique convergence is now considered to be an important part of collisional orogenic systems. I suggest that north-directed movement of the Dakota block resulted in transpression along the east side of the Black Hills. Evidence for transpressional tectonic motion, presented below, is based on:
101 1. The presence of the major structural discontinuity that is present in the study area northeast of the Harney Peak Granite that truncates earlier structures.
2. The north-northwest continuation of this structural discontinuity.
3. The ability of transpression to explain lithostratigraphic differences between fault-blocks and the previously recognized “cryptic” folding event across the central and northern Black Hills.
4. The ability of transpression to explain the intrusion, distribution, and metamorphism of the Harney Peak Granite.
5. The paucity of pre-1715 Ma porphyroblasts northeast of the structural discontinuity.
6. The bimodal distribution of 40Ar/39Ar dates on either side of the structural discontinuity.
7. The overall pattern of faulting in the east-central and northeastern Black Hills.
Structural Discontinuity
As discussed above, a major structural discontinuity exists along the
Keystone Fault over much of the study area. Along the west-central portion of
the study area, the structural discontinuity is marked by the Empire Mine Fault as displacement is transferred from the Keystone Fault to the Empire Mine Fault as they merge along a single fault-segment in the Hill City quadrangle (Ratte and
Wayland, 1969) and then diverge farther to the northwest (Figure 27). The structural discontinuity is then marked by the trace of the Empire Mine Fault to the northwest. The Empire Mine Fault continues northwest through the Hill City quadrangle (Ratte and Wayland, 1969) and the Silver City quadrangle (Redden et al., 2005). In the Silver City quadrangle, the Empire Mine Fault apparently
102 bifurcates into east and west branches (Redden et al., 2005). The structural discontinuity is presumably carried along the east branch of the Empire Mine
Fault as the west branch appears to tip out into a large fold nose (cf., Redden et al., 2005), although displacement may be transferred to the Silver City Fault.
The Silver City Fault is consistent with the biotite metamorphic isograd along most of its length (cf. Figure 6; Nabelek et al., in press) implying a metamorphic break across the fault zone.
Evidence noted above demonstrates that this discontinuity is a sinistral strike-slip fault zone, which predates the Harney Peak Granite intrusion. The discontinuity continues northwest through the Hill City and Silver City quadrangles where it truncates all earlier structures to the southwest (cf. Figure
7; Dewitt et al., 1989; Ratte and Wayland, 1969; Redden et al., 2005). The overall pattern of mapped faults northeast of the discontinuity (Figure 39; cf.
Figure 7) implies a dominantly strike-slip component to faulting; hence, the area northeast of the mapped discontinuity could represent various “tectonic slivers” moved unknown distances along the margin of the Archean Wyoming craton
(Figure 39; cf. Figure 27). The idea of the fault-bounded blocks representing
“tectonic slivers” is supported by the lack of determinate stratigraphic relations between fault-bounded blocks within the Mount Rushmore quadrangle.
Dissimilar structural histories for fault-blocks A, B, and C as compared to fault- blocks D – I and similar but unique structural histories for fault-blocks D – I also support this idea (Figures 5 and 15 – 20).
103
Figure 39 - Generalized geologic map of the northeastern Black Hills. Purple box outlines the Mt. Rushmore quadrangle and contains a scaled inset of detailed mapping completed during this study. Compare with Figures 7 and 27. Detailed mapping is needed to delineate the structural detail of the eastern Black Hills and elaborate on the proposed transpressional model.
Emplacement of the Harney Peak Granite
The mechanisms of ascent and emplacement of granitic magma are still an area of intense investigation and considerable debate (e.g., Paterson et al.,
1989; Simpson, 1999; Paterson and Schmidt, 1999; Brown and Solar, 1998 a,1998b). Granitic magmas are the principal agents of heat and mass transfer through the Earth’s crust (Yoshinobu et al., 1998 and references therein) and
104 have been extensively used to constrain the time of deformational events (e.g.,
Miller et al., 1999).
Understanding the timing and emplacement of plutons is critical to
interpretations of the age and significance of structures and metamorphism in
orogenic terranes. The spatial and temporal association of syntectonic granites
and regional tectonic structures has been demonstrated for many orogenic belts
(e.g., Weinberg, 1996; Brown and Solar, 1999), yet the mechanisms responsible
for ascent and emplacement of plutons are not well constrained. Ascent of
magmas is generally described by one of two end-member models: (1) diapirism
or (2) diking. Each model has its own inconsistencies in its ability to successfully
emplace substantial amounts of new material into the crust. The most basic
problem for both models is how room is made for the ascending magma bodies.
In convergent settings, it is clear that that there must exist structural
controls on the ascent and emplacement of plutons. Hence, it follows that there must be a relationship between deformational and thermal processes. Simpson
(1999) recognized that a valid argument could be made for the uniqueness of
each pluton with respect to its source, transport, and emplacement
characteristics. Brown and Solar (1998a, 1998b, 1999) suggested that in
transpressive orogens ascent of magma is controlled by tectonic structures.
Space is made both by the magma itself (ballooning, stoping, viscous flow of the wall rocks) and by creation of dilational space (e.g., Hutton, 1988; Glasner,
1991). Collins and Sawyer (1996) suggested a pervasive flow model for melt migration through partially molten crust. In this model, a feedback relationship
105 exists between deformation and melt generation. Melt pressure increase leads
to melt-enhanced embrittlement of the crust (Davidson et al., 1994). Shear
zones develop in response to rheologic heterogeneities in the crust that can be
enhanced by the presence of melt. The melt then migrates through the crust
along the shear zone extending the zone of pervasive flow (Brown and Solar,
1998a). Flow is driven by pressure gradients generated by tectonic stresses
over short distances and by magma buoyancy over longer distances (Rutter,
1997). Whatever the mechanisms for ascent of felsic magma, space must still be
made for their emplacement. Magmas clearly find or open space for themselves.
The main body of the Harney Peak Granite and its satellite plutons occupy
an area of roughly 100 km2 in the south-central Black Hills (Figure 27). The
mechanisms of emplacement of the main Harney Peak Granite has been inferred
by analogy from the study of one of its smaller satellites (~1 km2), the Calamity
Peak pluton (Duke et al., 1988). The model suggests that the Harney Peak
Granite consists of multiple pulses of genetically related magma that were initially
disconcordant sills fed by concordant feeder dikes. Each successive pulse was
injected forcefully along the upper contact between pre-existing granite layers
and the country rock. Continued growth by forceful injection resulted in
ballooning of the granite and upward and outward growth of the granite and produced an overall dome-shaped structure of the Harney Peak Granite.
I suggest that sinistral transpression along northwest directed strike-slip
faults (i.e. the structural discontinuity along the Keystone and Empire Mine faults)
would result in opening dilitational space that resulted in a low pressure gradient
106 allowing for the emplacement of magma. Nabelek et al. (in press) argue that no
doming actually took place during the emplacement of the Harney Peak Granite and that the fault-block containing the granite was uplifted along the Empire
Mine-Keystone faults on the north and the Hill City Fault along the west. They
cited the strong difference in foliation outside the Harney Peak block (Hill et al.,
2004; Ratte and Wayland, 1969) and metamorphic overprint of the Harney Peak
Granite across the structural discontinuity, along with the presence of late
andalusite as evidence of uplift. Movement along faults in the study area during
or after the emplacement of the Harney Peak Granite is demonstrated by
precipitation of graphite along fault zones (Nabelek et al., in press), prevalence of
secondary fluid inclusions within pegmatites that occur near faults (Sirbescu and
Nabelek, 2003), and overgrowth of M2 porphyroblasts in fault-related zones
(Figure 26).
If the overall motion along the Empire Mine-Keystone fault system was left
lateral, then a left-step over or releasing bend would have opened dilatational
space along a negative flower structure during transpression (Figures 40 and
41). The emplacement of the Harney Peak Granite would have been facilitated
by opening of dilitational space by northwest-directed sinistral transpression.
Generation of melt would have been accelerated by the release of confining
pressure and because feedback relationships exist between deformation and
melt generation (Brown and Solar, 1998a). Generation, segregation, and
emplacement of the granitic melt would have important affects on the thermal
and rheological behavior of the crust. The melt would have migrated through the
107 crust along an inferred listric shear zone (Nabelek et al., 2001) extending the
zone of pervasive flow (Brown and Solar, 1998a). Flow is driven by pressure
gradients generated by tectonic stresses over short distances and
Figure 40 - Diagram showing map view of mechanics of a releasing bend stepover along a sinistral strike-slip fault. Extension and normal faulting occur at high angles to the bounding strike-slip faults and are shown as hatched lines. Modified from Van der Pluijm and Marshak (2004).
by magma buoyancy over longer distances (Rutter, 1997). If deformation and
melt generation are coeval, the introduction of melt into the shear zone will
effectively weaken the crust and locally concentrate strain (Solar, 2006).
The presence of melt within the Harney Peak block would have made it
more buoyant than surrounding blocks (Nabelek et al., in press), resulting in uplift
that may have also been facilitated by inversion of the negative flower structure
during progressively more east-west directed compression as shown in Figures
15 - 20. This scenario is supported, in part, by the fact that the Harney Peak
108 Granite is spatially associated with the structural lowest (and presumably oldest)
Paleoproterozoic rocks of the younger sedimentary succession in the Black
Figure 41 - Block diagram of a negative flower structure. These structures may form from transtensile zones associated with releasing bends in an overall transpressional regime. Modified from Van der Pluijm and Marshak (2004).
Hills (Holm et al., 1997). The recumbent folds (F1/F2) recognized in this study
and elsewhere within the Harney Peak block would then be preserved D2a structures and as a result, F2a folds. These folds then record a weak S2a fabric.
Transposition of these early D2a features occurred outside the Harney Peak
block. Within the Harney Peak block, strain was concentrated by reaction-
enhanced ductility and melting, effectively partitioning the strain into the melt.
109 The lack of structural overprint by the Harney Peak Granite in adjacent fault-
blocks, even though there is a strong metamorphic overprint (Nabelek et al., in
press), may be explained by strain partitioning in the melt in the Harney Peak block such that little strain was transmitted to the country rocks of adjacent fault- blocks.
40 39 M1 Porphyroblasts and Ar/ Ar Data
One of the original goals of this study was to relate porphyroblast growth
to tectonic fabrics in order to form a basis for dating deformation fabrics. No M1 porphyroblasts were recognized in the study area, although spotted schists in fault-block B may preserve vestiges of M1 porphyroblasts (Figure 11b). Nabelek
et al. (in press) suggested that near the Harney Peak Granite M1 garnets were
wholly replaced by M2 garnets in response to elevated temperatures and fluid
flow. However, remnant M1 porphyroblasts are preserved to the west of the
defined structural break recognized in this study and in places affected by high-
grade M2 metamorphism. Moreover, porphyroblasts related to the thermal overprint of the Harney Peak Granite are ubiquitous in the study area indicating
that the rocks (on the east side of the structural break) had the correct chemistry to form early M1 porphyroblasts, yet no M1 porphyroblasts were found. Two
possibilities exist to explain this delimma: (1) the rocks of the study area initially
lacked the correct chemistry to form early porphyroblasts and metasomatic
reactions related to the Harney Peak Granite intrusion sufficiently altered the
chemistry of the rocks allowing M2 porphyroblasts to form; or (2) the rocks east of
110 the structural break initially had the correct chemistry to form early
porphyroblasts but were not subjected to the same P-T conditions as rocks that
preserve M1 porphyroblasts west of the structural discontinuity. Although
evidence is inconclusive, I believe that the lack of M1 porphyroblasts east of the
structural break is related to translation of the fault-bounded blocks as tectonic
slices along the margin of the Wyoming Province.
Available 40Ar/39Ar data (Dahl et al., 1999a) also show a remarkable trend
when one considers the structural discontinuity. The data show a bimodal age
distribution, with one group, SW of the discontinuity, having an average age of
~1735 Ma and a second group, NW of the discontinuity having an average age of
~1350 Ma (Table 4). Dahl et al. (1999a) stated that the cause of the bimodal ages of mica suites was uncertain but suggested that resetting by Tertiary intrusions was likely. When these data are considered with respect to the structural discontinuity defined during this study, all older dates are from samples west of the structural break and all younger dates are from samples east of the
discontinuity. 40Ar/39Ar data from Holm et al. (1997) also demonstrate that rocks
west of the structural break yield older dates. Data presented for muscovite and
biotite, (Dahl et al., 1999a) show broad flat plateaus, which are not consistent
with thermal resetting. Moreover, micas from the Nemo area (Figure 3) show a
broad range of dates (> 150 m.y.; Dahl et al., 1999a). This area is intensely
faulted and, although not conclusive, these age data demonstrate the need for
more investigation of the fault-bounded blocks as potentially separate tectonic
blocks with unique structural histories.
111 40Ar/39Ar data for hornblende (Dahl et al., 1999 a, b) show a classic “stair- case” pattern of increasing plateau ages. Thus ages determined by Dahl et al.
(1999 a, b) for hornblende are minimum estimates for the ages. All hornblende data were collected west of the structural break leaving no room for comparison of the same type data from the east-side of the structural break. Because no hornblende data east of the structural break are currently available for comparison, the overall pattern may indicate a normal cooling pattern across unrelated terranes. However, hornblende data are highly concordant with mica data from the west-side of the structural break, again supporting the disparity in age relationships across the structural discontinuity.
SW of the Structural Discontinuity NE of the Structural Discontinuity
Sample Mineral Age (Ma) Sample Mineral Age (Ma) BH-3 Hornblende 1691 ± 5 LEG-1 Biotite 1341 ± 4
CM-2242 Hornblende 1787 ± 8 LEG-8 Biotite 1307 ± 7 1736 ± T93-6 Hornblende LEG-10 Biotite 1396 ± 5 13 1714 ± T93-6 Hornblende LEG-15 Biotite 1286 ± 7 10 T93-6 Hornblende 1742 ± 6 LEG-25 Biotite 1457 ± 7
T93-6 Hornblende 1742 ± 5 NM-1 Muscovite 1361 ± 13
HM-1007** Muscovite 1221 ± 9 NM-3 Biotite 1308 ± 8
HM-32** Biotite 1304 ± 4
Table 4 - Summary of 40Ar/39Ar results from the northern Black Hills, Dahl et al. (1999a). **Italicized data are from gold-bearing quartz veins of the Homestake Mine, which lies along the projection of the structural discontinuity.
112 Overall Structural Pattern and “Cross Folds”
The overall structural pattern of faults east of the defined structural
boundary also strongly suggests strike-slip movement along the traces of
mapped faults (Figure 39). A lack of detailed mapping to the north of the Mount
Rushmore quadrangle limits this discussion to overall trends. Nevertheless, the
pattern is that of a series of anastamosing strike-slip faults and subsidiary faults.
Interpretation is complicated by the heterogeneity of the crust and preexisting
planar weaknesses (e.g. joints, old faults, foliations) that cause stress
concentrations and local changes in stress trajectories (Van der Pluijm and
Marshak, 2004). Gross lithologies are depicted in Figure 39, which was adapted
from the 1:250,000-scale map of the Black Hills (Dewitt et al., 1989). Although
some offset is indicated by the distribution of the rock types, there is no way of
knowing exactly how far any of these blocks may have been translated during
transpression from current evidence, nor have I considered any form of rigid
block rotation. In addition, we cannot assume that the present day erosional
surface is parallel to the σ1-σ3 plane of the strike-slip faults.
Transpression can also explain previously recognized but enigmatic
structural relationships in the Proterozoic core of the Black Hills, for example the
presence of “cross-folds” (cf. Redden et al., 1990; Noble et al., 1949), the
differences in lithostratigraphy among fault-bounded blocks in the east-central
Black Hills, and the chemical and structural dissimilarities that exist between the
two Archean granite exposures (Little Elk Mountain and Bear Mountain granites).
The F3 cross-folds, discussed previously, can be explained as en echelon arrays
113 of folds typically associated with continental strike-slip faults. It would be
advantageous to find and better describe these cross-folds and their relationship
to fault zones throughout the east-central Black Hills. This study presents evidence for dissimilar lithostratiphic relationships within the fault-blocks of the
Mount Rushmore quadrangle, which may be explained by transpression of these fault-blocks along the eastern margin of the Wyoming Province. Another example of the discontinuity of lithostratigraphy is the fact that the older
Paleoproterozoic succession is restricted to the Nemo area. Moreover, the mafic rocks of the Little Elk terrane are calc-alkaline, suggesting an arc-related genesis, while the mafic rocks of the Mount Rushmore quadrangle and elsewhere in the Black Hills have N-MORB affinities (Nabelek, unpublished data).
This strongly implies juxtaposition of terranes by transpression in the east-central
Black Hills. Much debate has occurred as to the dissimilarities (I-type versus S- type; REEs etc.) of the Little Elk and Bear Mountain Archean exposures, which is easily explained by being genetically unrelated and juxtaposed by transpression.
Application of the Transpressional Model to the Black Hills
A transpressional model for the Black Hills is a very attractive tectonic
paradigm as oblique convergence is now generally considered part of collisional
orogens. Previous tectonic models have called upon multiple sialic rifts in the
Black Hills to explain the lithologic differences juxtaposed over relatively small
distances (cf. Redden et al., 1990). Transpression can also explain the
differences in lithostratigraphy between fault-blocks. Previous workers have also
114 called upon cryptic tectonic events to explain various structures in the Black Hills,
for example a regional F1 folding event that produced no penetrative fabric or
cryptic tectonism to explain “cross-folds” (cf. Redden et al., 1990).
The transpressional model simplifies the overall tectonic setting with
respect to the number of events needed to produce the recognized deformation.
With a transpressional model, early north-directed folds (D2a) and later north-west
directed (D2b) folds could be products of the same transpressional event as the convergence direction changes. The emplacement of the Harney Peak Granite and the lack of agreement between the metamorphic and structural overprints related to emplacement can be explained with a transpressional model. The overall mechanism for the emplacement of the Harney Peak Granite can be explained in the context of available thermo-barometric studies (cf. Nabelek et al., in press) using a north-directed transpressional model. The late regional event resulting in a spaced cleavage over much of the southern Black Hills may also be related to final collision.
Transpressional Model
Earlier studies of the Black Hills called for as many as five separate deformation events and rather complex explanations for the tectonic history of the Black Hills and, by extrapolation, for southern Laurentia. This study considers three deformational events, discussed below, as a basis for formulating a transpressional tectonic model for the Black Hills and its relationship to the assembly of southern Laurentia.
115 Deformation in the Black Hills was produced by the collision of the
Wyoming Province and the Dakota block (e.g. Dahl et al., 1999 a, b;
Chamberlain et al., 2002). The deformation was partitioned into two events: (1)
the D2a event (correlative with the Black Hills orogeny), and (2) the D2b event
(correlative with the Dakotan orogeny). Figure 42 illustrates the possible
configuration of the Wyoming Province and Dakota block from 1780 – 1720 Ma.
Northwest directed translation of the Dakota block was a result of the 1780-1740
Ma Medicine Bow orogeny (Chamberlain, 1998). Oblique convergence ~1760 –
1740 Ma resulted in NW-SE compression (D2a) that formed the major early D2a folds (F1 nappe structures as reported in earlier studies: e.g., Redden et al.,
1990; Dahl et al., 1999 a, b). It is possible that the shape of the intervening
margins of the Wyoming Province and Dakota block played a substantial role
during the early D2a deformation, as the interaction of recesses and salients
produced complex interactions during their convergence. Collision between the
Wyoming Province and the Dakota block became progressively more west-
directed as shown by Figures 15 – 20 and likely coincides with the ~1755 – 1710
Ma ages of garnets from the Bear Mountain area (Dahl and Frei, 1998).
Continued northwest translation of the Dakota block (relative to the Wyoming
province as shown in Figure 42), resulted in the third period of deformation that
produced progressively more strike-slip movement along the eastern edge of the
Black Hills. Dilational space was opened up along the defined structural
boundary allowing for the emplacement of the Harney Peak Granite and related
deformation. Pieces of the intervening deformational zone between the
116 Wyoming Province and the Dakota block were then moved northwestward as the
Dakota block slipped past the eastern edge of the Wyoming Province.
Translation of these tectonic slivers resulted in F3 cross-folds that locally
overprint the dominant F2 folds. The late spaced cleavage recognized in this
study is interpreted to be a result of the subsequent Central Plains orogeny.
Figure 42 - Possible configuration of the Wyoming Province and Dakota block during 1760-1715 Ma collision. Red area denotes present day Black Hills uplift, Blue areas are other Laramide basement uplifts that expose Precambrian cores. Green area denoted the region affected by deformation.
Extrapolation to the Tectonics of the Wyoming Province
Mueller et al. (2005) presented a very good model for the assembly of
southern Laurentia invoking simultaneous SW-NE convergence of the Hearne,
Medicine Hat, and Wyoming provinces. However, they could not reconcile the
117 younger orogenic events that occurred along the eastern margin of the Wyoming
province (Black Hills and Dakota orogens). Nor did they take into account the
recently recognized correlation of the Dakota block with the the Sask craton (e.g.,
Bickford et al., 2005) which underlies the Glennie and Flinn Flon Domains of the
Trans-Hudson.
One of the major shortcomings of the Mueller et al. (2005) model is that it requires the Black Hills to be part of the 1.97 – 1.85 Ga Trans-Hudson orogeny based on superseded dates for tectonothermal events in the Black Hills as reported by Gosselin et al. (1988) and Redden et al. (1990). There is no evidence that the Black Hills underwent collisional tectonism prior to 1790 Ma
(e.g., Dahl et al., 1998; Dahl et al., 1999 a, b; Chamberlain et al., 2002;
McCombs et al., 2003). The Mueller et al. (2005) model, however, improves
upon the model of Dahl et al. (1999 a, b) which required two orthogonal
subduction zones, one dipping westward underneath the Wyoming Province and
one dipping northward underneath the Wyoming and Superior Provinces.
The assembly of southern Laurentia involved collisions between the
Wyoming Province, the Superior craton, the Medicine Hat block, the Sask craton,
and the Hearne-Rae Province that occurred prior to the circa 1600 Ma Central
Plains orogeny. Collectively, the number of terranes and the temporal overlap of
their juxtapositions present many geodynamic problems. Figure 43, which is an
attempt to address some of the major problems, illustrates a possible
configuration for major Archean and Paleoproterozoic crustal blocks incorporated
into southern Laurentia between 1.9 and 1.7 Ga.
118 I suggest that convergence of the Medicine Hat block , Hearne province,
and Wyoming province occurred much as depicted by Mueller et al. (2005).
However, convergence along the subduction zone shown in Figure 43 was not a
steady state. I propose that subduction between the Superior and Hearne-Rae
provinces was fast, while the western portions of the zone it was slower
Figure 43 - Depiction of potential configuration of major cratonic blocks of now part of southern Laurentia. Black arrows show direction of movement. Subduction zone polarity shown by teeth that point in direction of subduction. DB = Dakota Block.
(i.e. Medicine Hat-Hearne, Wyoming-Medicine Hat). This resulted in a large-
scale tear fault orthogonal to the north-directed subduction as shown in Figure
43. The Trans-Hudson proper was a result of fast, north-dipping subduction that
juxtaposed the Hearne-Rae and Superior cratons and accreted intervening
terranes. This fast subduction led to the Sask craton being overthrusted by the
arc-related rocks of the Glennie and Flinn-Flon domains (Chiarenzelli, 2002) as
first suggested by Lewery et al. (1990). In the west, collision between the
Medicine Hat and Hearne provinces resulted in the Vulcan tectonic zone, but
119 timing of this collision remains uncertain. Subsequent collision between the
Medicine Hat block and the Wyoming Province resulted in the Great Falls
Tectonic Zone, which is at least temporally equivalent to the Trans-Hudson.
During this collision, northwest directed translation of the faster moving eastern
side of the subduction zone brought the Dakota block into the intervening ocean
between the Wyoming and Superior Provinces. By ~1820 – 1775 Ma, the north-
directed subduction that sutured the Hearne and Superior provinces had become
a left-lateral strike-slip boundary (Annesley et al., 2005; White et al., 1999;
Lithoprobe Project, http://www.litho.ucalgary.ca/atlas/thot/thot _blurb.html). This left lateral strike-slip led to the terminal collision between the Wyoming and
Superior provinces during the Black Hills and Dakotan orogenies.
CONCLUSIONS
1. Three deformational events are recognized in the Mt. Rushmore
quadrangle. Table 3 summarizes these events and retains the
numbering system of previous studies to avoid unnecessary confusion.
The earliest event incorporates both the D1 and D2 events of previous
studies. Previous studies ascribed the D1 event to a nebulous north-
directed collision and the D2 event to the terminal collision between the
Wyoming and Superior Archean provinces. I have chosen to call the
earliest event recognized in the study area D2 and separate it into D2a and
D2b subevents. I do so, knowing that it is likely the earliest event that
120 affected all the Paleoproterozoic rocks in the Black Hills, but have chosen
to keep the D2 moniker to denote that both subevents are related to the
Wyoming-Superior collision. An oblique collision of the intervening Dakota
block during terminal Wyoming-Dakota block-Superior collision resulted in
portioning the D2 event into the D2a and D2b subevents. The D2a subevent
formed the earliest folds and faults recognized in the Black Hills and
convergence was more north directed and resulted in early north directed
compression. During oblique collision, the overall convergence became
more northwest directed transpression, resulting in the strong overprint by
D2b folding and fabrics as the Dakota block slipped past the eastern
Wyoming margin. The D4 event is the second event to affect the
Paleoproterozoic rocks of the study area and is related to the
emplacement of the Harney Peak Granite. Major deformation associated
with the D4 event in the study area is restricted to fault-block A. The D5
event affects the southern half of the study area and is likely a result of
far-field stresses related to the Central Plains orogen.
2. The study area contains nine fault bounded blocks with separate
structural histories. The differences in structural patterns, deformation,
and metamorphism may be explained by a transpression model.
3. A major structural discontinuity exists along the Keystone and
Empire Mine Faults in the Black Hills, South Dakota. Evidence from
121 detailed mapping in the Mount Rushmore quadrangle demonstrates that
this discontinuity is a sinistral strike-slip fault zone. This discontinuity
continues northwest through the Hill City and Silver City quadrangles
where it truncates all earlier structures to the southwest. The overall
pattern of mapped faults north east of the discontinuity implies a sinistral
transpressional event; hence, the area northeast of the mapped
discontinuity represent various “tectonic slivers” moved unknown
distances along the margin of the Archean Wyoming craton.
4. Emplacement of the Harney Peak Granite occurred during the
waning stages of D2b as dilational space was created in a releasing
bend during north directed transpression. Uplift of the fault-bounded
block containing the Harney Peak Granite is likely a result of differential
buoyancy related to the emplacement of the melt (Nabelek et al., in press)
coupled with late compression of the releasing bend reactivating the faults
in high-angle reverse motion.
5. Movement along strike-slip faults continued after the emplacement
of the Harney Peak Granite. Mapped faults within the study area are
characterized by brittle deformation. Evidence for post-intrusion
movement include: (1) abundance of graphite along the major fault zones
(Nabelek et al., in press); (2) intense metamorphism by the Harney Peak
Granite and presence of late andalusite recording polybaric uplift of the
122 Harney Peak block coupled with lack of intense deformation outside the
Harney Peak block; and (3) overgrowth of Harney Peak related
porphyroblasts in and near strike-slip faults mapped during this study.
Furthermore, faults mapped during this study truncate structures
interpreted to be directly related to the Harney Peak intrusion and
therefore must be somewhat younger in age. It is postulated that the
overall sinistral transpression is related to the collision of the Cheyenne
Belt from the south prior to the terminal east-west collision of the Wyoming
and Superior cratons. Similar timing of collisional events in the Laramie
Range and Hartville Uplift in Wyoming supports this hypothesis.
6. The structural pattern observed in the study area, lack of correlative
stratigraphy between fault-bounded blocks in the study, and overall
structure of the east-central Black Hills may be best explained by an
oblique convergence model.
7. The Dakota segment represents a younger orogenic event related to
the assembly of southern Laurentia. The deformation recorded in the
Paleoproterozoic rocks of the Laramide Black Hills uplift is not related to
the Trans-Hudson orogen. The age of deformational events, paucity of
intervening juvenile terranes, and presence of the Dakota block preclude
the southern continuation of the Trans-Hudson orogen.
123 124 REFERENCES
Alabi, A.O., Camfield, P.A., and Gough, D.I., 1975, The North American Central Plains anomaly. Geophysical Journal of the Royal Astronomical Society, v. 43, p. 815-834.
Annesley, I.R., Madore,C., and Portella, P., 2005, Geology and thermotectonic evolution of the western margin of the Trans-Hudson Orogen; evidence from the eastern sub-Athabasca basement, Saskatchewan: Canadian Journal of Earth Sciences, v. 42, n. 4, pp. 573-597.
Baird, D.J., Knapp, J.H., Steer, D.N., Brown, L.D., Nelson, K.D., 1995, Upper-mantle reflectivity beneath the Williston basin, phase-change Moho, and the origin of intracratonic basins: Geology, v. 23, n. 5, p. 431-434.
Baird, D. J., Nelson, K. D., Knapp, J. H., Walters, J. J., and Brown, L. D., 1996, Crustal structure and evolution of the Trans-Hudson orogen: Results from seismic reflection profiling: Tectonics, v. 15, p. 416-426.
Balk, R., 1931, Inclusions and foliation of the Harney Peak granite, Black Hills, South Dakota, Jour. Geol., vol. 39, p. 736-748.
Bayley, R.W., 1970, Structure and Mineralization of Precambrian rocks in the Galena-Roubaix district, Black Hills, South Dakota: U.S. Geological Survey, Bulletin 1312-E, 15 p.
Bayley, R. W., 1972a, Geologic field compilation map of the northern Black Hills, South Dakota: United States Geological Survey Open-File Report 72-29, scale 1:48,000.
Bayley, R. W., 1972b, Preliminary geologic map of the Nemo district, Black Hills, South Dakota: United States Geological Survey Miscellaneous Investigation Series Map I-712, scale 1:24,000.
Bayley, R. W., 1972c, Preliminary report on the geology and gold deposits of the Rochford district, Black Hills, South Dakota: United States Geological Survey Bulletin 1332-A, 24 p.
Bekker, A., Eriksson, K.A., 2003, A Paleoproterozoic drowned carbonate platform on the southeastern margin of the Wyoming Craton; a record of the Kenorland breakup: Precambrian Research, v. 120, issue 3-4, p. 327-364.
Bickford, M. E., Van Schmus, W. R., and Zeitz, I., 1986, Proterozoic history of the mid-continent region of North America: Geology, v. 14, p. 492-496.
Bickford, M.E., Collerson, K.D., Lewry, J.F., Van Schmus, W.R., and Chiarenzelli, J.R., 1990, Proterozoic collisional tectonism in the Trans-Hudson orogen, Saskatchewan: Geology, v. 18., p. 14-18.
Bickford, M. E., Mock, T. D., Steinhart, W. E., III 1, Collerson, K. D., and Lewry, J. F., 2005, Origin of the Archean Sask Craton and its Extent Within the Trans-Hudson Orogen:
125 Evidence from Pb and Nd Isotopic Compositions of Basement Rocks and Post-Orogenic Intrusions: Canadian Journal of Earth Sciences, v. 42, p. 659-684.
Bird, P., 1998 Kinematic history of the Laramide orogeny in latitudes 35-49 N, western United States: Tectonics, v. 17, p. 780-801.
Boerner, D.E., Kurtz, R.D., and Craven, J.A., 1996, Electrical conductivity and Paloproterozoic foredeeps. Journal of Geophysical Research, v. 101, p. 13 775 – 13 789.
Boerner, D.E., Craven, J.A., Kurtz, R.D., Ross, G.M., and Jones, F.W., 1998, The Great Falls tectonic zone: suture or intracontinental shear zone?: Canadian Journal of Earth Sciences, v. 35, p. 175-183.
Boone, E. and Hynes, A., 1990, A structural cross-section of the northern Labrador Trough, New Quebec in The early Proterozoic Trans-Hudson Orogen of North America, Lewry, J.F. and Stauffer, M.R., editors: Geologicial Association of Canada Special Paper, v. 37, pp. 387 – 396.
Brown, M., and Solar, G.S., 1998a, Shear zone systems and melts: feedback relations and self- organization in orogenic belts: Journal of Structural Geology, v. 20, p. 211-227.
Brown, M., and Solar, G.S., 1998b, Granite ascent and emplacement during contractional deformation in convergent orogens: Journal of Structural Geology, v. 20, p. 1365-1393.
Brown, M., and Solar, G.S., 1999, The mechanism of ascent and emplacement of granite magma during transpression; a syntectonic granite paradigm: Tectonophysics, v. 312, issue 1, p. 1-33.
Camfield, P.A., and Gough, D.I., 1975, Anomalies in daily variation magnetic fields and structure under north-western United States and south-western Canada. Geophysical Journal of the Royal Astronomical Society, v. 41, p. 193-218.
Camfield, P.A., and Gough, D.I., 1977, A possible Proterozoic plate boundary in North America, Canadian Journal of Earth Science, v. 14, p. 1229-1238.
Camfield, P.A., Gough, D.I., and Porath, H., 1970, Preliminary results from geomagnetic deep sounding studies in the northwestern U. S. and southwest Canada: Eos, Transactions, American Geophysical Union, Vol. 51, Issue 4, pp. 268
Chamberlain, K. R., 1998, Medicine Bow orogeny: Timing of deformation and model of crustal structure produced during continent-arc collision, ca. 1.78 Ga, southeastern Wyoming: Rocky Mountain Geology, v. 33, p. 259-277.
Chamberlain, K.R., Frost, C.D., and Frost, B.R., 2003, Early Archean to Mesoproterozoic evolution of the Wyoming Province; Archean origins to modern lithospheric architecture: Canadian Journal of Earth Sciences = Revue Canadienne des Sciences de la Terre, Vol. 40, Issue 10, pp. 1357-1374
126 Chamberlain, K. R., Bauer, R. L., Frost, B. R., and Frost, C. D., 2002, Dakotan orogen: Continuation of Trans-Hudson orogen or younger, separate suturing of Wyoming and Superior cratons. Geological Association of Canada – Mineralogical Association of Canada Abstracts Volume, 27, 18.
Chiarenzelli, J.R., and Roden-Tice, M., 2002, History and Tectonic evolution of the Sask Craton, Trans-Hudson orogen: Geologic Association of Canada Abstracts with Programs.
Collins, W.J. and Sawyer, E.W., 1996, Pervasive granitoid mamga transfer through the lower- middle crust during non-coaxial compressional deformation: Journal of Metamorphic Geology, v. 14, pp. 565-579.
Condie, K.C., and Shadel, C.A., 1984, An early Proterozoic volcanic arc succession in southeastern Wyoming: Canadian Journal of Earth Sciences, v. 21, p. 415-427.
Connolly, J.P. and O’Harra, C.C., 1929, The mineral wealth of the Black Hills: South Dakota School of Mines and Technology Bulletin, v. 16, 418p.
Dahl, P.S. and Frei, R., 1998, Single-phase dating of coexisting garnet and saurolite in the Black Hills collisioinal orogen (South Dakota), with implications for Early Proterozoic tectonism: U.S. Geology, v. 26, p. 111-114.
Dahl, P. S., Frei, R., and Dorais, M. J., 1998, When did the Wyoming Province collide with Laurentia?: New clues from step-leach Pb-Pb dating of garnet independent of its inclusions: Geological Society of America Abstracts with Programs, v. 30, no. 7, p. A109.
Dahl, P. S., Foland, K. A., Gardner, E. T., Holm, D. K., and Hubacher, F.A., 1999a, New constraints on the timing of Early Proterozoic tectonism in the Black Hills (South Dakota), with implications for docking of the Wyoming province to Laurentia: Geologic Society of America Bulletin, v. 111, p. 1335-1349.
Dahl, P. S., Dorais, M. J., Roberts, H. J., Kelley, S. P., and Frei, R., 1999b, Electron microprobe geochronometry of age-zoned monazite crystals in Archean metapelites from the Wyoming Province; the nature of Pb rejuvenation and implications for regional tectonism, Geological Society of America, 1999 annual meeting: Boulder, CO, United States, Geological Society of America (GSA), p. 39.
Dahl, P. S., Hamilton, M., Stern, R., Frei, R., and Berg, R., 2000, In situ SHRIMP investigation of an Early Proterozoic metapelite, with implications for Pb-Pb step leach dating of garnet and staurolite. Geological Society of America Abstract Program 32(7):A297.
Dahl, P. S., Hamilton, M. A., Wooden, J. L., and Frei, R., 2003, Evidence for 2480 Ma rifting in the Black Hills, S. Dakota: U-Pb ages of sphene and zircon from the Blue Draw metagabbro sil, and their tectonic significance. Geological Society of America Abstracts with Programs, 35, 506.
Dahl, P.S., Terry, M.P., Jercinovic, M.J., Williams, M.L., Hamilton, M.A., Foland, K.A., Clemet, S.M., and Friberg, L.M., 2005a, Electron probe (Ultrachron) microchronometry of
127 metamorphic monazite: Unraveling the timing of polyphase thermotectonism in the easternmost Wyoming Craton (Black Hills, South Dakota): American Mineralogist, v. 90, p. 1712-1728
Dahl, P. S., Hamilton, M. A., Jercinovic, M. J., Terry, M. P., Williams, M. L., and Frei, R., 2005b, Comparative isotopic and chemical geochronometry of monazite, with implications for U- Th-Pb dating by electron microprobe: An example from metamorphic rocks of the eastern Wyoming craton (U.S.A.). American Mineralogist, 90, 619-638.
Darton, N. H. and Paige, S., 1925, Central Black Hills [quadrangle], South Dakota: United States Geological Survey Geologic Atlas of the United States, Folio 107, 34 p.
Davidson, C., Schmid, S.M., and Hollister, L.S.,1994, Role of melt during deformation in the deep crust: Terra Nova, v. 6, p. 133-142.
DeWitt, E., Redden, J.A., Wilson, A.B., Busher, D., and Dersch, J.S., 1986, Mineral resources potential and geology of the Black Hills National Forest, South Dakota and Wyoming: United States Geologic Survey Bulletin 1580, 135 p.
DeWitt, E., Redden, J.A., Wilson, A.B., and Buscher, D., 1989, Geologic map of the Black Hills area, South Dakota and Wyoming: United States Geological Survey Miscellaneous Investigation Series Map I-1910, 1 sheet, scale 1:250,000.
Dheeradilok, P., 1974, A detailed study of the deformational history of part of the Rockerville area, Black Hills, South Dakota: Unpublished Masters Thesis, South Dakota School of Mines and Technology.
Distlehorst, J., 1999, Analysis of Regional Metamorphic and Deformational Fabrics of the Southern Black Hills, South Dakota; Unpublished Master’s Thesis, University of Missouri-Columbia, 84 p.
Duebendorfer, E.M., and Houston, R.S., 1987, Proterozoic accretionary tectonics at the southern margin of the Archean Wyoming craton: Geological Society of America Bulletin, v. 98, p. 554-568.
Duke, E.F., Redden, J.A., and Papike, J.J., 1988, Calamity Peak layered granite-pegmatite complex, Black Hills, South Dakota: Part I. Structure and emplacement: Geological Society of America Bulletin, v. 100, p. 825-840.
Duncan, D.W.,1982, Kinematic and structural analysis of pre-Beltian rocks, west-central Tobacco Root Mountains, Southewest Montana: Rocky Mountain section of Geologicial Society of America, Abstracts with Programs, v 14, pp. 310.
Duncan, I.J., 1984, Structural evolution of the Thor-Odin gneiss dome: Tectonophysics, v. 101, pp. 87-130.
Dutch, S.I., 1983, Proterozoic structural provinces in the north-central United States, Geology, v. 11, p. 478-481.
128 Ellis, M.A., 1986, Lithospheric strength in compression; subduction initiation, migration of orogeny, and an origin of exotic terranes: Abstracts with Programs - Geological Society of America, Vol. 18, Issue 6, pp. 594
Evans, S., Jones, A.G., Spratt, J., and Katsube, J., 2003, Central Baffin electromagnetic experiment (CBEX): Phase 2, Geological Survey of Canada Current Research, 2003- C24: 10.
Evans, S., Jones, A.G., Spratt, J., and Katsube, J., 2005, Central Baffin electromagnetic experiment (CBEX) maps the NACP in the Canadian arctic, Physics of the Earth and Planetary Interiors, v. 150, p. 227-237.
Faure, G., 1986, Principles of isotope geology: New York, John Wiley and Sons, 589 p.
Ferguson, I.J., Stevens, K.M., and Jones, A.J., 2005, Electrical-resistivity imaging of the central Trans-Hudson Orogen in eastern Saskatchewan, Canada, Canadian Journal of Earth Sciences, v. 42: this issue.
Fisher, D.J., 1942, Preliminary report on some pegmatites of the Custer district: South Dakota Geological Survey Report Investigations, v. 44, 35 p.
Fisher, D.J., 1945, Preliminary report on the mineralogy of some pegmatites near Custer: South Dakota Geological Survey Report Investigations, v. 50, 92 p.
Frost, B.R., 2000, Late Archean structural and metamorphic history of the Wind River Range; evidence for a long-lived active margin on the Archean Wyoming Craton: Geological Society of America Bulletin, Vol. 112, Issue 4, pp. 564-578.
Frost, C.D., and Chamberlain, K.R., 1998, Archean evolution of the Wyoming Province: evidence from the Big Horn Mountains: Geological Society of America Bulletin, Vol. 30, Issue 7, pp. 395.
Garcia, X., and Jones, A.G., 2005, Electromagnetic image of the Trans-Hudson orogen – THO94 transect, Canadian Journal of Earth Sciences, v. 42: this issue.
Gardner, E.T., Dahl, P.S., Holm, D.K., and Foland, K.A., 1996, (super 40) Ar/ (super 39) Ar thermochronology of micas from the Little Elk Granite and environs, northern Black Hills, South Dakota: Geological Society of America Abstracts with Programs, v. 28, issue 4, p. 9.
Glasner, A.F., 1991, Plutonism, oblique subduction and continental growth: an example from the Mesozoic of California: Geology, v. 19, p. 784-786.
Goldich, S.S., Lidiak,E.G.,Hedge,C.E., and Walthall, F.G., 1966, Geochronology of the mid- continent region, United States 2. Northrern area: Journal of Geophysical Research, v. 71, p. 5389-5408.
Gosselin, D.C., Papike, J.J., Zartman, R.E., Peterman, Z.E., and Laul, J.C., 1988, Archean rocks of the Black Hills, South Dakota: Reworked basement from the southern extension
129 of the Trans-Hudson orogen: Geological Society of America Bulletin, v. 100, p. 1244- 1259.
Gosselin, D.C., Papike, J.J., Shearer, C.K., Peterman, Z.E., and Laul. J.C., 1990, Geochemistry and origin of Archean granites from the Black Hills, South Dakota, Canadian Journal of Earth Sciences, v. 27, p. 57-71.
Green, A.G., Cumming, G.L., and Cedarwell, D., 1979, Extension of the Superior-Churchill boundary zone into southern Canada: Canadian Journal of Earth Sciences, v. 16, p. 1691-1701.
Green, A.G., Stephenson, O.G., Mann, G.D., Kanasewich, E.R., Cumming, G.L., Hajnal, Z., Mair, J.A., West, G.F., 1980, Cooperative seismic surveys across the Superior-Churchill boundary zone in southern Canada: Canadian Journal of Earth Sciences, v. 17, issue 5, p. 617-632.
Green, A.G., Hajnal, Z., and Weber, W., 1985, An evolutionary model for the western Churchill Province and western margin of the Superior Province in Canada and the north-central United States: Tectonophysics, v. 116, p. 281-322.
Gwynne, C.S., 1944, Pegmatites in the Beecher Rock basin: South Dakota Geological Survey Report Investigations, v. 48.
Hatcher, R.D., Jr., and Williams, R.T., 1986, Mechanical model for single thrust sheets; Part I, Taxonomy of crystalline thrust sheets and their relationships to the mechanical behavior of orogenic belts: Geological Society of America Bulletin, v. 97, issue 8, p. 975-985.
Helms, T.S., and Labotka, T.C., 1991, Petrogenesis of Early Proterozoic pelitic schists of the southern Black Hills, South Dakota: Constraints on regional low-pressure metamorphism: Geological Society of America Bulletin, v. 103, p. 1324-1334.
Herman, J.D., Harder, A.S., and Roberts, D.V., 1986, Tracing Precambrian lithologies and structures beneath northeastern part of Michigan Basin, AAPG Eastern Section meeting; abstracts: Tulsa, OK, United States, American Association of Petroleum Geologists, p. 1067.
Hess, F.L., 1925, The natural history of pegmatites: Engineering Mining Journal-Press, v. 120, p. 289-298.
Hill, J.C., Nabelek, P.I., and Bauer, R.L., 2004, Differential deformational history of fault- bounded blocks: ‘southern Trans-Hudson’ Orogen, Black Hills South Dakota: Geological Society of America Abstracts with Programs, v. 36, pp. 569.
Hobbs, B.E., Means, W.D., and Williams, P.F., 1976, An Outline of Structural Geology: New York, John Wiley and Sons, 571 p.
Hoffman, P.F., 1981, Autopsy of Athapuscow aulacogen: a failed arm affected by three collisions, in Proterozoic basins of Canada, ed. Campbell, F.H.A., Geological Survey of Canada, Paper 81-10, p. 97-102.
130
Hoffman, P.F., 1987, Early Proterozoic foredeeps, foredeep magmatism, and Superior-type iron-formations of the Canadian Shield, in Proterozoic Lithospheric Evolution, ed. Kroner, A., American Geophysical Union, Geodynamics Series 17, p. 85-98.
Hoffman, P.F., 1988, United Plates of America: the birth of a craton, Ann. Rev. Earth Planet. Sci., 16, 543-603.
Hoffman, P.F., 1989, Precambrian geology and tectonic history of North America, in The geology of North America – an overview, eds. Bally, A.W., and Palmer, A.R., Geological Society of America, Boulder, CO, p. 447-511.
Hoffman, P.F., 1990, Subdivision of the Churchill Province and extent of the Trans-Hudson orogen, in The Early Proterozoic Trans-Hudson Orogen of North America, eds. Lewry, J.F., Stauffer, M.R., Geological Association of Canada Special Paper 37, p. 15-39.
Holm, D.K., Dahl, P.S., and Lux, D.R., 1997, 40Ar/39Ar evidence for Middle Proterozoic (1300- 1500 Ma) slow cooling of the southern Black Hills, South Dakota midcontinent, North America: Implications for Early Proterozoic P-T evolution and posttectonic magmatism: Tectonics, v. 16, p. 609-622.
Holm, D.K., Darrah, K.S., and Lux, D.R., 1998, Evidence for widespread ~1760 Ma metamorphism and rapid crustal stabilization of the Early Proterozoic (1870-1820 Ma) Penokean orogen, Minnesota: American Journal of Science, v. 298, p. 60-81.
Houston, R.S., 1993, Late Archean and Early Proterozoic geology of southeastern Wyoming, in Snoke, A.W., Steidtmann, J.R., and Roberts, S.M., eds., Geology of Wyoming: Geological Survey of Wyoming Memoir 5, p. 78-116.
Houston, R.S., Duebendorfer, E.M., Karlstrom, K.E., and Premo, W.R., 1989, A review of the structure of the Cheyenne belt and Proterozoic rocks of southern Wyoming, in Grambling, J.A., and Tewksbury, B.J., eds., Proterozoic geology of the southern Rocky Mountains: Geological Society of America Special Paper 235, p. 1-12.
Hutton, D.H.W., 1988, Granite emplacement mechanisms and tectonic controls; inferences from deformational studies, in The origin of granites, Brown, P.E. chair, Transactions of the Royal Society of Edinburgh: Earth Sciences, v. 79, Issue 2-3, p. 245-255.
Jones, A.G., and Savage, J.G., 1986, North American conductivity anomaly goes east. Geophysical Research Letters, v. 13, p. 685-688.
Jones, A.G., Craven, G.A., McNeice, G.A., Ferguson, I.J., Boyce, T., Farquharson, C., and Ellis, R.G., 1993, The North American Central Plains conductivity anomaly within the Trans- Hudson orogen in northern Saskatchewan, Geology, v. 21, p. 1027-1030.
Jones, A.G., Katsube, T.J., and Schwann, P., 1997, The longest conductivity in the world explained: sulphides in fold hinges causing very high electrical anisotropy, Journal of Geomagnetism and Geoelectricity, v. 49, p. 1619-1629.
131 Jones, A.G., Ledo, J., and Ferguson, I.J., 2005, Electromagnetic images of the Trans-Hudson orogen: the North American Central Plains (NACP) anomaly revealed, Canadian Journal of Earth Sciences, v. 42(4): this issue.
Karlstrom, K.E., and Houston, R.S., 1984, The Cheyenne Belt: Analysis of a Proterozoic suture in southern Wyoming: Precambrian Research, v. 25, p. 415-446.
Karlstrom, K.E. and CD-Rom Working Group, 2002, Structure and Evolution of the Lithosphere beneath the Rocky Mountains: Initial Results from the CD-ROM experiment: GSA Today, v. 12, no. 3.
Klasner, J.S., and King, E.R., 1986, Precambrian basement geology of North America and South Dakota, Canadian Journal of Earth Science, v. 23, p. 1083-1102.
Klasner, J.S., and King, E.R., 1990, A model of tectonic evolution of the Trans-Hudson orogen in North and South Dakota, in The Early Proterozoic Trans-Hudson orogen of North America, pp. 271-286, Geol. Assoc. of Can., St. Johns, NF.
Landes, K.K., 1928, Sequence of mineralization in the Keystone, South Dakota, pegmatites: American Mineralogist, v. 13, issue 10, p. 519-530.
Lanphere, M.A., Wasserburg, G.J.F., Albee, A.L., and Tilton, G.R., 1964, Redistribution of strontium and rubidium isotopes during metamorphism, World Beater Complex, Panamint Range, California, in Craig, H., Miller, S.L., and Wasserburg, G.J.F., eds., Isotopic and cosmic chemistry: North Holland Publications, p. 269-320.
Latham, T.S., Best, J., Chaimov, T.A., Oliver, J., Brown, L., Kaufman, S., 1988, COCORP profiles from the Montana plains; the Archean cratonic crust and a lower crustal anomaly beneath the Williston Basin: Geology Boulder, v. 16, issue 12, p. 1073-1076.
Laubscher, H., and Bernoulli, D., 1982, History and deformation of the Alps, in Hsü, K., ed., Mountain building processes: London, Academic Press, p. 169-180.
Lewry, J.F., and Collerson,K.D., 1990, The Trans Hudson Orogen: extent, subdivision, and problems, in The Early Proterozoic Trans Hudson Orogen of North America, eds. Lewry, J.F., and Stauffer, M.R., Geological Association of Canada, Special Paper 37, p. 1-14.
Lewry, J.F., Sibbald, T.I.I., and Schledewitz, D.C.P., 1985, Variation in character of Archean rocks in the western Churchill province and its significance, in Ayres, L.H., Thurston, P.C., Card, K.D., and Weber, W., eds., Evolution of Archean supracrustal sequences: Geological Association of Canada Special Paper 28, p. 239-261.
Lewry, J.F., Thomas, D.J., MacDonald, R., and Chiarenzelli, J., 1990, Structural relations in accreted terranes of the Trans-Hudson orogen, Saskatchewan; telescoping in a collisional regime? In The Early Trans-Hudson Orogen of North America, eds. Lewry, J.F., Stauffer, M.R., Geological Association of Canada, Special Paper 37, p. 75-94.
Lewry, J.F., Hajnal, Z., Green, A., Lucas, S.B., White, D., Stauffer, M.R., Ashton, K.E., Weber, W., and Clowes, R., 1994, Structure of a Paleoproterozoic continent-continent collision
132 zone: a LITHOPROBE seismic reflection profile across the Trans-Hudson Orogen, Canada, Tectonophysics, v. 232, p. 143-160.
Lisenbee, A.L., 1978, Laramide structure of the Black Hills Uplift, South Dakota-Wyoming- Montana, in Laramide folding associated with basement block faulting in the western United States, ed. Matthews, V., III: Geological Society of America Memoir, issue 151, p. 165-196.
Lisenbee, A.L., and Dewitt, E., 1993, Laramide evolution of the Black Hills uplift, in Snoke, A.W., Steidtmann, J.R., and Roberts, S.M., eds., Geology of Wyoming: Geological Survey of Wyoming Memoir 5, p. 374-412.
Lucas, S.B., Green, A., Hajnal, Z., White, D., Lewry, J., Ashton, K., Weber, W., and Clowes, R., 1993, Deep seismic profile across a Proterozoic collision zone: surprises at depth, Nature, v. 363, p. 339-342.
Lucas, S.B., White, D., Hajinal, Z., Lewry, J., Green, A., Clowes, R., Zwanzig, H., Ashton, K., Schledewitz, D., Stauffer, M., Norman, A., Williams, P.F., Spence, G, 1994, Three- dimensional collisional structure of the Trans-Hudson Orogen, Canada, in Clowes, R.M., and Green, A.G. ed, Seismic reflection probing of the continents and their margins, Tectonophysics, v. 232, issue 1-4, p. 161-177.
Machado, N., 1990, Timing of collisional events in the Trans-Hudson Orogen; evidence from U- Pb geochronology for the New Quebec Orogen, the Thompson Belt, and the Reindeer Zone (Manitoba and Saskatchewan), in Lewry, J.F., and Stauffer, M.R., eds., The early Proterozoic Trans-Hudson Orogen of North America, Geological Association of Canada Special Paper, v. 37, p. 433-441.
McCombs, J.A., Dahl, P.S., and Hamilton, M.A., 2003, U-Pb ages of Neoarchean granitoids from the Black Hills, South Dakota, USA: implications for crustal evolution in the Archean Wyoming province: Precambrian Research, v. 130, p. 161-184.
McCombs, J.A., Dahl, P.S., and Hamilton, M.A., 2004, Ion microprobe study of zircon geochronology in basement granitoids from the Black Hills, South Dakota, USA, with implications for evolution of the Archean Wyoming Province, western Laurentia. Precambrian Research, 130, 161-184.
Miller, B.V., Samson, S.D., and D’Lemos, R.S., 1999, Time span of plutonism, fabric development, and cooling in a Neoproterozoic magmatic arc segment: U-Pb age constraints from syn-tectonic plutons, Sark, Channel Islands, U.K.: Tectonophysics 312, p. 79-95.
Mueller, P.A., Wooden, J.L., Nutman, A.P., and Mogk, D.W., 1988, Early Archean crust in the northern Wyoming Province; evidence from U-Pb ages of detrital zircons: Precambrian Research, v. 91, issue 3-4, p. 295-307.
Mueller, P.A. Hetherington, A.L., Kelly, D., Wooden, J., and Mogk, D., 2002, Paleoproterozoic crust within the Great Falls tectonic zone: implications for the assembly of southern Laurentia: Geology 20: p. 127 – 130.
133
Mueller, P.A., Burger, R., Wooden, J.L., Brady, J.B., Cheney, J.T. Harms, T.A., Heatherington, A.L., and Mogk, D.W., 2005, Paleoproterozoic Metamorphism in the Northern Wyoming Province: Implications for the Assembly of Laurentia: The Journal of Geology, v. 113, p. 169-179.
Mukhopadhyay, M., and Gibb, R.A., 1981, Gravity anomalies and deep structure of eastern Hudson Bay: Tectonophysics, v. 72, issue 1-2, p. 43-60.
Nabelek, P. I., Russ-Nabelek, C., and Denison, J. R., 1992a, The generation and crystallization conditions of the Proterozoic Harney Peak leucogranite, Black Hills, South Dakota, USA: Petrogeologic and geochemical constraints: Contributions to Mineralogy and Petrology, v. 110, p. 173-191.
Nabelek, P.I., Russ-Nabelek, C., and Haeussler, G.T., 1992b, Stable isotope eveidence for the petrogenisis and fluid evolution in the Proterozoic Harney Peak leucogranite, Black Hills, South Dakota., Geochimica et Cosmochimica Acta, v. 56, p. 403-417.
Nabelek, P.I., Liu, M., and Sirbescu, M., 2001, Thermo-rheological, shear heating model for leucogranite generation, metamorphism, and deformation during the Proterozoic Trans- Hudson orogeny, Black Hills, South Dakota. Tectonophysics, 342, 371-388.
Nabelek, P.I., Huff, T.A., Wilke, M., 2002, Carbonic fluid production during regional and contact metamorphism in the Black Hills, USA: Geochimica et Cosmochimica Acta, v. 66, issue 15A, p. 543.
Nabelek, P.I., Labotka, T.C., Helms, T., and Wilke, M., in press, Fluid-Mediated polymetamorphism related to Proterozoic collision of Archean Wyoming and Superior provinces in the Black Hills, South Dakota:
Nelson, K.D., Baird, D.J., Walters, J.J., Hauck, M., Brown, L.D., Olivier, J.E., Ahern, J.L., Hajnal, Z., Jones, A.G., and Sloss, L.L., 1993, Trans-Hudson orogen and Williston basin and North Dakota: new COCORP deep profiling results, Geology, v. 21, p. 447-450.
Noble, J.A., and Harder, J.O., 1948, Stratigraphy and metamorphism in a part of the northern Black Hills and the Homestake Mine, Lead, South Dakota, Geol. Soc. Am., Gull., vol. 59, p. 941-976.
Noble, J.A., Harder, J.O., and Slaughter, A.L., 1949, Structure of a part of the northern Black Hills and the Homestake mine, Lead, South Dakota: Geological Society of America Bulletin, v. 60, p. 321-352.
Norton, J.J., 1974, Gold in the Black Hills, South Dakota, and how new deposits might be found: U. S. Geological Survey Circular, U. S. Geological Survey, Reston, VA, 22 p.
Norton, J.J., 1976, Field compilation map of the geology of the Keystone area, Black Hills, South Dakota, USGS open file report, 76-0297.
134 Norton, J.J., and Redden, J.A., 1990, Relations of zoned pegmatites to other pegmatites, granite, and metamorphic rocks in the southern Black Hills, South Dakota: American Mineralogist, v. 75, p. 631-655.
Norton, J.J., Page, L.R., and Brobst, D.A., 1962, Geology of the Hugo pegmatite, Keystone, South Dakota: U.S. Geological Survey Professional Paper 297-B, p. 49-126.
Paterson, S.R., and Schmidt, K.L., 1999, Is there a close spatial relationship between faults and plutons?: Journal of Structural Geology, v. 21, issue 8-9, p. 1131-1142.
Paterson, S.R., Vernon, R.H., Tobisch, O.T., 1989, A review of criteria for the identification of magmatic and tectonic foliations in granitoids, Journal of Structural Geology, v. 11, p. 349-363.
Peterman, Z.E., Sims, P.K., Zartman, R.E., Schulz, K.J., 1985, Middle Proterozoic uplift events in the Dunbar Dome of northeastern Wisconsin, USA: Contributions to Mineralogy and Petrology, v. 91, issue 2, p. 138-150, Springer International: Heidelberg-New York.
Pignotta, G.S., and Benn, K., 1999, Magnetic fabric of the Barrington Passage pluton, Meguma Terrane, Nova Scotia: a two-stage fabric history of syntectonic emplacement, Tectonophysics, v. 307, p. 75-92.
Premo, W.R., and Van Schmus, W.R., 1989, Zircon geochronology of Precambrian rocks in southeastern Wyoming and northern Colorado, in Grambling, J.A., and Tewksbury, B.J., eds., Proterozoic geology of the southern Rocky Mountains: Geological Society of America Special Paper 235, p. 13-48.
Ramsay, J.G., 1967, Folding and Fracturing of Rocks: New York, McGraw-Hill, 568 p.
Ratte, J.C., 1986, Geologic map of the Medicine Mountain quadrangle, Pennington county, South Dakota: United States Geological Survey Miscellaneous Investigation Series Map I-1654, scale 1:24,000.
Ratte, J.C., and Wayland, R.G., 1969, Geology of the Hill City quadrangle, Pennington county, South Dakota: A preliminary report: United States Geologic Survey Bulletin 1271-B, 14 p.
Redden, J.A., 1963, Geology and pegmatites of the Fourmile quadrangle, Black Hills, South Dakota: U.S. Geological Survey Professional Paper, 297-D, p. 199-291.
Redden, J.A., 1968, Geology of the Berne quadrangle, Black Hills, South Dakota: United States Geological Survey Professional Paper 297-F, p. 343-408.
Redden, J.A., 1981, Summary of the geology of the Nemo area, in Rich, F. J., ed., Geology of the Black Hills, South Dakota and Wyoming, Geological Society of America Guidebook, Rocky Mountain Section, Annual Meeting, Rapid City, South Dakota, 1981: American Geological Institute, Falls Church, VA, p. 193-210.
135 Redden, J.A., and Dewitt, E., 1996, Early Proterozoic tectonic history of the Black Hills – An atypical Trans-Hudson orogen: Geological Society of America Abstracts with Programs, v. 28, no. 7, p. A-315.
Redden, J.A., and Duke, E.F., 1996, Proterozoic tectonic/plutonic development and associated metamorphism/metasomatism in the southern Black Hills: Road Log, Field Trip 9, South Dakota School of Mines Bulletin, 19, 78-89.
Redden, J.A., and Lisenbee, A.L., 1996, Geologic setting, Black Hills, South Dakota in Paterson, C.J., and Kirchner, J.G., eds., Guidebook to the geology of the Black Hills, South Dakota: Rapid City, South Dakota School of Mines and Technology Bulletin 19, p. 1-8.
Redden, J.A., Peterman, Z.E., Zartman, R.E., and DeWitt, E., 1990, U-Th-Pb geochronology and preliminary interpretation of Precambrian tectonic events in the Black Hills, South Dakota, in Lewry, J.F., and Stauffer, M.R., eds., The Trans-Hudson orogen: Geologcial Association of Canada Special Paper 37, p. 229-251.
Redden, J.A., Alexander, D., and Nonnast, D., 2005, Geologic Map of the Silver City Quadrangle, South Dakota: Department of Environment and Natural Resources Geologic Quadrangle map 6, 1:24000 scale.
Rutter, E. H., 1997, The influence of deformation on the extraction of crustal melts: a consideration of the role of melt-assisted granular flow: In Holness, M. (ed.), Deformation-Enhanced Melt Segregation and Metamorphic Fluid Transport: The Mineralogical Society Series 8, London: Chapman and Hall, p. 82-110.
Shackelton, R.M., and Ries, A.C., 1984, The relation between regionally consistent stretching lineations and plate motions: Journal of Structural Geology, v. 6, p. 111-120.
Sheridan, V.M., Stephens, H.G., Staatz, M.H., and Norton, J.J., 1957, Geology and beryl deposits of the Peerless pegmatite, Keystone, South Dakota: U.S. Geological Survey Professional Paper 297-A, p. 1-47.
Simpson, C., 1999, Introduction: ‘The influence of granite emplacement on tectonics’: Tectonophysics, v. 312, p. vii-viii.
Sims, P.K., Peterman, Z.E.,1986, Early Proterozoic Central Plains orogen: A major buried structure in the north-central United States, Geology, 14, p. 488-491.
Sims, P.K., Van Schmus, W.R., Schultz, K.J., Peterman, Z.E., 1989, Tectono-stratigraphic evolution of the Early Proterozoic Wisconsin magmatic terranes of the Penokean Orogen, Canadian Journal of Earth Science, v. 26, p. 2145-2158.
Sims, P.K., Peterman, Z.E., Hildebrand, T.G., and Mahan, S., 1991, Precambrian basement map of the Trans-Hudson orogen and adjacent terranes, northern Great Plains, U. S. A.: United States Geological Survey Miscellaneous Investigations Series Map I-2215, 53 p., 1 sheet, scale 1:250,000.
136 Sims, P.K., Schulz, K.J., DeWitt, E., Brasaemle, B., 1993, Petrography and geochemistry of early Proterozoic granitoid rocks in Wisconsin magmatic terranes of Penokean Orogen, northern Wisconsin: U. S. Geological Survey Bulletin, p. J1-J31.
Sirbescu, M.C. and Nabelek, P.I., 2003, Evolution of magmatic fluids in the Harney Peak Granite and associated pegmatites, Black Hills, South Dakota: Evidence from fluid inclusions: Geochimica et Cosmochimica Acta, v. 67, n. 13., p. 2443-2465.
Solar, G.S., 2006, Migmatites, granites, and shear zones: the signature of Appalachian sutures: Geological Society of America Abstracts with Programs, Northeastern Section, v. 38, n.2, p. 87.
Stauffer, M.R., 1984, Manikewan; an early Proterozoic ocean in central Canada, its igneous history and orogenic closure: Precambrian Research, v. 25, issue 1-3, p. 257-281.
Terry, M.P., and Friberg, L.M., 1990, Pressure-temperature-time path related to the thermotectonic evolution of an Early Proterozoic metamorphic terrane, Black Hills, South Dakota: Geology, v. 18, p. 786-789.
Thomas, M.D., 2001, Potential field images of the Proterozoic Trans-Hudson orogen: in Canadian Shield: Significance for mapping buried extensions in plate tectonic models: American Geophysical Union Abstracts with Programs, Spring Meeting.
Toft, P.B., Scowen, P.A.H., Arkani-Hamed, J., Francis, D., 1993, Demagnetization by hydration in deep-crustal rocks in the Grenville Province of Quebec, Canada; implications for magnetic anomalies of continental collision zones: Geology Boulder, v. 21, issue 11, p. 999-1002.
Van Hise, C.R., 1890, The pre-Cambrian rocks of the Black Hills, Geol. Soc. Am., Bull., vol. 1, p. 203-243.
Van Schmus, W.R., Bickford, M.E., Lewry, J.F., Macdonald, R., 1987, U-Pb geochronology in the Trans-Hudson orogen, northern Saskatchewan, Can. J. of Earth Sci., 24, 407-424.
Van Wyck, N., and Johnson, C.M., 1997, Common lead, Sm-Nd, and U-Pb constraints on petrogenesis, crustal architecture, and tectonic setting of the Penokean Orogeny (Paleoproterozoic) in Wisconsin: Geological Society of America Bulletin, v. 109, issue 7, p. 799-808.
Van der Pluijm, B.A., and Marshak, S., 2004, Earth Structure: An introduction to structural geology and tectonics, 2nd ed.: New York, W. W. Norton & Company, 656.
Vannay, J.C. and Steck, A., 1995, Tectonic evolution of the High Himalaya in Upper Lahul (NW Himalaya, India): Tectonics, v. 14, n. 2, p. 253-263.
Walker, R.J., Hanson, G.N., Papike, J.J., and O’Neil, J.R., 1986, Nd, O and Sr isotope constraints on the origin of Precambrian rocks, southern Black Hills, South Dakota: Geochimica et Cosmochimica Acta, v. 50, p. 2833-2846.
137 Weinberg, R. T., 1996, Ascent mechanisms of felsic magmas: news and views: Royal Society of Edinburgh Transactions, Earth Science, v 87, p. 95-103.
White, D.J., Lucas, S.B., Hajnal, Z., Green, A.G., Lewry, J.F., Weber, W., Bailes, A.H., Syme, E.C., and Ashton, K., 1994, Paleo-Proterozoic thick-skinned tectonics: LITHOPROBE seismic reflection results from the eastern Trans-Hudson Orogen, Canadian Journal of Earth Sciences, v. 31, p. 458-469.
White, D.C., Helmstaedt, H.H., Harrap, R.M., Thurston, P.C., van der Velden, A., Hall, K., Davis, D.W., 1999, Accretionary tectonics in the late Archean? First results from the Lithoprobe Western Superior Transect: Lithoprobe Report, University of British Columbia, Lithoprobe Secretariat for the Canadian Lithoprobe Program, Vancouver, BC, Canada, 168 p.
Wilson, J.T., 1962, The structure and origin of continents: ICSU Rev., v. 4, issue 4, p. 205-215.
Wooden, J., and Mueller,P., 1988, Pb, Sr, and Nd isotopic compositions of a suite of Late Archean, igneous rocks, eastern Beartooth Mountains: Implications for crust-mantle evolution: Earth and Planetary Science Letters, v. 87, p. 59-72.
Woodland, B.G., 1979, Geometry and origin of deformational structures in the Precambrian metamorphic rocks of the Hill City area, Black Hills, South Dakota: Contributions to Geology, v. 17, no. 1, p. 1-23.
Yoshinobu, A. S., Okaya, D.A., and Patterson, S.R., 1998, Modeling the thermal evolution of fault-controlled magma emplacement models: implications for the solidification of granitoid plutons: Journal of Structural Geology, v. 20, No. 9/10, p. 1205-1218.
Zartman, R.E., and Stern, T.W., 1967, Isotopic age and geologic relationships of the Little Elk Granite, northern Black Hills, South Dakota: U.S. Geological Survey Professional Paper 575-D, p. D157-D163.
Zartman, R.E., 1986, Age of the granite at Bear Mountain and the Bear Mountain Dome, in Ratte, J.C., 1986, Geologic map of the Medicine Mountain quadrangle, Pennington County, South Dakota: U.S. Geological Survey Miscellaneous Geologic Investigation Map 1-1654, scale 1:24,000.
138 APPENDIX A – Lithostratigraphic Units
The following is a list of mappable units used during this study and corresponds to the
units on Plate I. This list is not a stratigraphy and implies no stratigraphic or age
relationships beyond the fact that the Harney Peak Granite is the youngest rock in the
study area.
Lithostraphic Units
XHPG - Harney Peak Granite and related pegmatites. Fine grained to pegmatitic, S-type
granite. Typically layered parallel or subparallel to dominant foliation. Dominant
minerals include oligoclase, microcline, perthite, quartz, muscovite, biotite, and
tourmaline. Mineralogy may vary with location, especially with respect to the
pegmatites.
Xqms – Quartz-mica schists. May contain minor amounts of staurolite and/or garnet with
sillimanite prevalent around the contact with the Harney Peak Granite. Dominant
minerals include quartz and biotite with lesser amounts of muscovite, opaques in
various amounts. Dominant biotite schists interlayered with subordinate
metagreywackes. May contain layers of quartzite.
Xqqm – Dominantly interlayered quartz schists, quartz-mica schists, phyllites, and mica
schists. Thin bedded. Locally shows good primary bedding features. Layers
may have relict “spots” of unknown porphyroblasts crossing beds at an angle and
defining the axial-plane of larger scale folds.
139 Xgs – Quartz-biotite-garnet schists, graphitic in part, interlayered with quartz-mica-garnet-
staurolite schists with quartz pebbles (matrix supported metaconglomerate)
Xsgw – Interlayered quartz-mica schists and metagreywacke, with subordinate graphitic
schists, slates, and phyllites. Similar to unit Xgs but with thicker-bedded
metagreywackes prominent with lesser banded quartzite. No porphyroblasts.
Xgw – Medium to thick bedded metagreywacke and quartzite; locally interlayered with
quartz-schists and minor phyllites which are common in the northwest part of the
quadrangle. Thick bedded quartzose rocks in the northeast part of the
quadrangle may have black, glassy beds of graphitic quartzite.
Porphyroblasts of garnet and/or staurolite common in more pelitic layers.
Xss – Staurolite-quartz-mica schist and biotite-staurolite-garnet schists with very large
staurolite porphyroblasts. Distinctive unit.
Xqgw – Mainly interlayered quartz-schist, quartz-mica schists ± garnet and/or staurolite
with interlayered metasiltstones and minor phyllites and slates. Layers of
greywacke are locally common. Similar to unit Xgw but with much less
metagreywacke and quartzose rocks.
Xs – Grey to black slates and “spotted” slates and schists with lenses of metasilttone.
140 Xtgw – Interlayered greywacke and slates, mica-garnet schists, and minor quartzite.
Similar to units Xgw and Xqgw but with more schistose rocks having garnet
porphyroblasts.
Xqs – quartz schists with minor quartz-biotite schists. Dominantly clean quartz-schists
with lenses of more pelitic mica schist. Porphyroblasts include staurolite and
garnet in the southeast of the quadrangle.
Xqps – Quartz phyllite and phyllonitic quartz-schist. Lower grade equivalent of unit Xqs.
Xbs – Biotite schist ± garnet.
Xamg – Amphibolite, amphibolite schist, metagabbro, and metabasalt undifferentiated.
Xbif – Iron formation and Fe-rich schists, cherty quartzite, and cummingtonite-grunerite
schists. Typically reddish-brown iron-stained beds. Typically shows sulfide
weathering.
Xqc – Quartzite and cherty quartzite. Medium-bedded, glassy in part.
Xtbd – Turbidite sequences. Thin to medium-bedded greywacke to clean quartz schists.
White to light grey in color.
Xaqm – Arkosic metagreywackes with interlayered quartz-schist and minor, incomplete
turbidite sequences. More micaceous but similar to unit Xtbd.
141 Xqbg – Quartz-biotite-garnet schists and metagreywacke. Graphitic in part. Similar to unit
Xsgw.
Xgph – Dark graphitic and sulfide rich schists and shales. Highly altered in many places
with residual graphite coating. May contain garnets and or andalusite.
142 APPENDIX B - Lithotectonic Units versus Stratigraphy
A major objective of this study was to provide information on stratigraphic relationships within the Mount Rushmore Quadrangle that would be applicable to ongoing and future investigations within the crystalline core of the Black Hills.
Published material concerning stratigraphic relationships within the Precambrian rocks is sparse and occurs mainly in non peer-reviewed sources (e.g. USGS
Professional Papers, various field guides). In addition, the number of workers historically and currently involved in such investigations is quite modest. Most contemporary investigations of the Precambrian geology of the Black Hills rely on three primary sources for stratigraphic interpretation, two of which predate the advent of plate tectonic theory: Noble and Harder (1948) and Redden (1963,
1968). Furthermore, available data regarding the age of most of the stratigraphic succession is temporally broad, dividing the rocks into two groups: (1) an older,
~2550 – 2480 Ma rift related sequence (Redden, 1981; Dahl et al., 2005a,b) that is unconformably overlain by (2) ~2015 – 1885 Ma (Bekker et al, 2003: Dahl et al., 2005a,b) continental shelf-rise, shallow-water deposits (Redden et al., 1990).
The difficulty in defining and correlating stratigraphic units with any confidence within the crystalline Black Hills is three-fold. First, the most basic problem is the recurrence of identical rock types within the succession and the paucity of good stratigraphic marker units. Second, the rocks within the crystalline core have been multiply and complexly deformed and great care must
143 be taken with regard to stratigraphic relationships. Thirdly, the rocks have
undergone at least two episodes of prograde metamorphism. The thermal
overprint of the Harney Peak granite has significantly altered the country rocks
within its aureole and the degree of alteration is not uniform.
Previous Work
Original stratigraphic work in the Precambrian Black Hills was carried out to evaluate the economic opportunity related to recently discovered gold deposits and mineralization associated with numerous pegmatite intrusions. The Lead area in the northern Black Hills, home of the Homestake mining company, served as a focus of investigation (e.g. Hosted and Wright, 1923; Paige, 1924;
McLaughlin, 1931; Gustafson, 1933). Noble and Harder (1948) presented a comprehensive study of the stratigraphy and metamorphism of the Lead area and defined a local stratigraphy. In the 1960’s, Redden (1963, 1968) mapped two 7.5-minute quadrangles in the southwestern Black Hills as part of the USGS study of pegmatites and other Precambrian rocks in the southern Black Hills, also defining a local stratigraphy. Subsequent workers in the southern Black Hills adopted the stratigraphic nomenclature of Redden.
The three primary sources mentioned above (Noble and Harder, 1948;
Redden, 1963, 1968) serve as the basis for almost all stratigraphic relationships cited in the subsequent literature for the Precambrian Black Hills. A fourth often cited source, Ratte and Wayland (1969), outlined stratigraphic relationships in the Hill City quadrangle, west of the Mount Rushmore Quadrangle, and
144 attempted to correlate rock units with the work of Redden (1963, 1968). For
reasons outlined below, this correlation is suspect, yet has remained the type
example for correlating stratigraphy across the Black Hills. Redden et al. (1990)
outlined a tentative correlation of Precambrian rock units across the Black Hills.
However, upon careful reading of the literature from which that correlation was
derived, several glaring inconsistencies were revealed, both in the actual
correlation and its use in interpreting the structure of the Black Hills. Figure 2-1
compares the various stratigraphic interpretations of the previously cited works
along with works from Bayley (1972b, 1972c).
Noble and Harder (1948) presented the first comprehensive study of the
stratigraphy and metamorphism of the Lead district in the northern Black Hills.
The stratigraphy as defined by Noble and Harder (1948) is shown in Figure 2-2.
No attempts to correlate the Lead stratigraphy with rocks in other parts of the
Black Hills were made with the exception of the Rochford district approximately
20 miles south of Lead, where the lowermost stratigraphy was repeated (Noble and Harder, 1948). This correlation was based on multiple lines of evidence, but primarily on the recognition of the Homestake formation and the ore mineralization that occurred within it. As stated by Noble and Harder (1948):
“There is not only a repetition of distinctive formations or members, but a repetition of a sequence of distinctive formations, and the repetition of an unconformity. It is fortunate that the evidence for correlation is so positive, since the whole stratigraphic column is upside-down.” Noble and Harder (1948) went on to say:
145 “We know of no other area in the Black Hills which can be surely correlated with the succession at Lead.”
Other authors (e.g. Runner, 1934) had suggested that the stratigraphic sequence
at Lead was repeated elsewhere in the northern Black Hills, but Noble and
Harder (1948) noted that these correlations were based on slight evidence and that the rock types included in the Lead district stratigraphy were not uncommon
“in all Precambrian rocks of the sedimentary (Algonkian) type.” They cautioned that correlation based just on similar rock types should not be made in the absence of the repetition of nearly all the sequence of the district.
Redden (1963) mapped the Four-mile quadrangle in the southwestern
Black Hills. Approximately one-half of the quadrangle is underlain by
Precambrian rocks, while the other half is overlain by Paleozoic and Tertiary strata. Redden defined three geologic formations in the Precambrian rocks of the Four-mile quadrangle. They are, in ascending order, the Bugtown formation, the Crow formation, and the Mayo formation (Figure 2-3a). Redden did not attempt to correlate the rocks of the Four-mile quadrangle with other rocks in the
Black Hills, but did note the lithologic similarity between the Precambrian rocks of the quadrangle and those defined by Noble and Harder (1948).
Redden (1968) mapped the Berne quadrangle immediately north of the
Four-mile quadrangle. The main structural element in the Berne quadrangle is the Grand Junction fault, which runs from the lower southeast side of the quadrangle slightly toward the northwest (Figure 2-3b). Redden separated
146
147
the rocks of the Berne quadrangle into two groups: (1) rocks to the west of the
Grand Junction fault; and, (2) those east of the Grand Junction fault. Rocks located west of the Grand Junction fault were divided into the three formations defined by Redden (1963) in the Four-mile quadrangle, and two previously undefined formations: the Vanderlehr formation and the Loues formation. The complete stratigraphy of the rocks located west of the Grand Junction fault as defined by Redden is shown in Figure 2-3b. Redden divided the rocks of the
148
149 Berne quadrangle located east of the Grand Junction fault into four separate
units as shown in Figure 2-3b. Redden stated that the stratigraphic relationship
of these four units was not fully known nor had any genetic relationship been established across the Grand Junction fault. However, he attempted to correlate the rock units on either side of the fault based solely upon lithologic similarity.
Ratte and Wayland (1969) divided the rocks of the Hill City quadrangle
(Figure 2-4) into two formations: the Bugtown formation as defined by Redden
(1963) and the previously undefined Oreville formation. The stratigraphy of the
Hill City quadrangle as defined by Ratte and Wayland (1969) is shown in (Figure
2-4). The adoption of the Bugtown formational name for metagraywackes, quartz-mica schists, and phyllites in the Hill City quadrangle was made seemingly without regard to the tenuous and somewhat speculative correlation of rock units across the Grand Junction fault in the Berne quadrangle. Ratte and Wayland
(1969) did not address this concern in their report. However, they did recognize the problem using their own Oreville formation in correlating rocks of similar lithology across the Hill City quadrangle. Most notably, they pointed out that rocks “of Oreville lithology” in the northeast corner of the quadrangle (across the
Empire Mine-Keystone fault system) were not necessarily correlative with
Oreville-type rocks elsewhere in quadrangle. Norton (1974) correlated the
Oreville formation with the Swede Gulch Formation of Bayley (1972) and then correlated the Swede Gulch with the Grizzly Formation of Noble and Harder
(1948).
150
Redden et al. (1990) proffered a first-order assessment of stratigraphic relationships for all of the crystalline rocks of the Black Hills. This correlation included the seminal papers mentioned above and also the work of Hosted and
Wright (1923), Bayley (1972, 1972b), and Redden (1981). The exact reasoning behind correlating the different stratigraphies across the Black Hills is not well explained, and seems in large part based on the unpublished work of Redden or master’s theses directed by Redden. Many of the assertions and correlations of
Redden et al. (1990) remain unclear and subject to much debate. Norton (1974) stated that neither during work on the Berne quadrangle (Redden, 1968), nor during subsequent mapping of adjacent areas, had a persuasive correlation been made between rocks on either side of the Grand Junction Fault. The tentative correlation of Redden et al. (1990) is reproduced in Table 2-1.
151
Problems with stratigraphic correlation
Stratigraphic relationships must exist within the metasedimentary sequences of rocks in the Black Hills. Previous studies (e.g. Noble and Harder,
1948; Redden, 1963, 1968) have demonstrated that sufficient evidence exists to establish stratigraphic relationships, at least locally, for parts of the crystalline
Black Hills. However, even the local stratigraphies are subject to certain ambiguities that are compounded when trying to correlate locally derived stratigraphic relationships from different parts of the Black Hills. Among the most fundamental problems are lack of distinct marker units and sequences, recognition of stratigraphic younging criteria, repetition of similar rock units, facies changes, and complex and incompletely understood structural relationships.
Noble and Harder (1948, 1949) recognized the perils of interpreting a stratigraphic succession in the Precambrian rocks of the Black Hills and using it to interpret structure. They cautioned that changes in their stratigraphic interpretation would necessitate a major revision of the structure. Ratte and
Wayland (1969) echoed this premise regarding the use of preserved primary structures and their interaction with tight minor folds as stratigraphic younging criteria. Locally and for brief distances in the field, relict primary stratigraphic structures such as graded bedding, load casts, and cross-beds can be used to determine the direction of stratigraphic younging. However, the complexity of
152 multiple fold generations and incomplete exposure can lead to erroneous assumptions without careful and consistent observation. Woodland (1979) also recognized difficulties in lithologic correlation with previously defined units and suggested that facies changes in lithologic members and soft-sediment deformation further exacerbated the potential for errors in structural interpretation based on proposed stratigraphic sequences. Furthermore, the occurrence of primary sedimentary structures is not consistent and varies considerably with lithology and even location.
Throughout most of the crystalline Black Hills, there is a paucity of traceable stratigraphic marker units, while there is an abundance of similar, indistinct rock types. Noble and Harder (1948) relied on the distinctive
Homestake formation to develop their stratigraphy. They described the majority rock type present in the other formations as gray phyllites and commented on the absence of appropriate marker units. Redden (1963) relied on the distinctive
Crow formation to delineate between the quartz-mica schists of his Bugtown formation and the lithologically similar Mayo formation. Ratte and Wayland
(1969) relied on gross lithology to separate the rocks in the Hill City quadrangle: thick bedded metagraywackes with inter-layered quartz-mica schists and phyllites of the “Bugtown formation” and the mica-schists of the Oreville formation. They also acknowledged the uncertainty in calling all similar sequences of mica-schist Oreville formation, most notably they singled out the rocks northeast of the Empire Mine fault. Even within the defined local stratigraphies, there may be multiple units that are so lithologically similar as to
153 be indistinguishable. Furthermore, distinct sequences of deposition are virtually
absent. Relative thickness of units and overall lithologic variation may be primarily controlled by subsequent deformation. For example, an inter-layered sequence of greywacke and shale may show a totally different ratio of greywacke to shale in the limbs of major folds as compared to the noses of the folds, giving the overall impression of two distinctive depositional regimes, while in fact only one existed.
Perhaps the most obvious demonstrations of the difficulty associated in correlating stratigraphy in the Precambrian Black Hills rocks is found in the geologic map of Redden et al. (1990; Figure 2-5). This map is a simplified version of Dewitt et al. (1989) 1:250,000-scale map of the Black Hills and has essentially the same description of map units. Table 2-1 reproduces the description of map units for the map of Redden et al. (1990). Some subtle but very enlightening discrepancies appear in the explanation as well as the tentative correlation of stratigraphies presented in Table 2-1. If one assumes the correlations of Redden et al. (1990) to be correct, the map explanation and hence the overall map must be incorrect with respect to the correlations and previous work. The following is a partial list of incorrect relationships derived from the descriptions of the map units and stratigraphic correlations as proposed by Redden et al. (1990; c.f. Figure 2-5, Table 2-1) based on comparison with the earlier published work (e.g. Noble and Harder, 1948; Redden, 1963, 1968; Ratte and Wayland, 1969):
154 1. Unit Xps includes part of the Oreville formation but overlies unit Xc which
is Crow formation. Ratte and Wayland (1969) clearly indicate that the
Oreville formation is older than the Bugtown formation. Redden clearly
indicates (1963, 1968) that the Bugtown formation is conformably overlain
by the Crow formation.
2. Unit Xps then cannot be the youngest metasedimentary unit. Unit Xqg
includes part of the Ellison formation (Noble and Harder, 1948), the top of
which is marked by an unconformity. This unit overlies unit Xbs of the
Oreville formation of Ratte and Wayland (1969). The Oreville formation
must occur above this unconformity based on Redden (1963, 1968) and
Ratte and Wayland (1969).
3. Unit Xgw breaks out the graywackes from four different stratigraphic units
that are not equivalent: Roubaix formation (Bayley, 1972b), Mayo and
Bugtown formations (Redden, 1968) and the Zimmer Ridge member of the
Oreville formation (Ratte and Wayland, 1969); and lumps them based on
similar rock type. The map pattern and structure resolved from this is
suspect, especially if one assumes the stratigraphy correlative. Either the
structure (as depicted) is correct and the stratigraphic relationships
incorrect, or vice versa. Both cannot be correct.
155
Figure 2-5. Reproduction of the simplified geologic map of Precambrian units in the Black Hills from Redden et al. (1990).
156 Table 2-1. Description of map units for Figure 2-5. Reproduced from Redden et al. (1990).
Ti Tertiary igneous rocks
Pr Phanerozoic rocks
Early Proterozoic Unconformity
Harney Peak granite. Coarse-grained to pegmatitic muscovite granite Layer metagabbro. Gravity differentiated Blue Draw Xh Xlg and pegmatite Metagabbro sill (2,700 Ma; Redden, 1981)
Xps Phyllite and slate. Mica schist in Lead area. Includes Grizzly Fm (Dodge, 1942), Swede Gulch Fm (Bayley, 1972), part of the lower Oreville Fm (Ratte and Wayland, 1969), and a large area of rocks in the core of the Black Hills
Biotite schist, calcsilicate gneiss, and amphibolite. Crow Fm (Redden, Metaconglomerate and quartzite. Includes the Xc Xcq 1963) wouthwest of Custer Boxelder Creek Fm, Benchmark Iron-Fm (Redden, 1981), and an unnamed carbonate-silicate-oxide facies Xgb Metagabbro. Sills and dikes of amphibolite or greenstone. At least two iron-formation along the west side of the Nemo area. separate ages (~1,883 Ma and 1,964 Ma) known and younger metagabbro intrudes the 1,884 Ma unit Xpb
Xpb Phyllite and biotite schist. Contains minor chert and amphibole- Unconformity bearing rocks. Locally intruded by thin sills of metagabbro. Includes Poverty Gulch slate, Nahant schist, and Irish Gulch slate in the Rochford area (Bayley, 1972) and Northwestern and Flagro
Muscovite schist and phyllite. Includes part of the Oreville Fm. Xms Archean Equivalent to unit Xpb in Rochford and Lead areas.
Xqg Quartzite and metagraywacke. Quartzite and siliceous schist. Includes Wgr Granite. Granite and gneissic granite. Includes Little Ellison Fm in the Lead area, Moonshine Gulch quartzite north of Elk Granite and the granite at Bear Mountain Rochford, and siliceous graywacke west of Rochford. Possibly equivalent in part to Xgw and Xcg in central Black Hills. U
Biotite schist or phyllite. Thin bedded and commonly garnet-rich Older metasedimentary rocks. Includes Nemo iron- Xbs Wos schist. Largely includes parts of the Oreville Formation formation and schist at Tinton area.
Xcg Conglomeratic biotite schist and phyllite. Also siliceous biotite phyllite, garnetiferous schist, quartite, and carbonate-silica iron-formation. Unconformable lower contact
Metabasalt near Rockerville. Includes metabasalt in Rochford, Lead, Xby and Rockerville areas
Xgw Metagraywacke. Siliceous mica schist and impure quartzite at repeated stratigraphic intervals. Includes part of Roubaix Fm (Bayley, 1972b), Mayo and Bugtown Fms (Redden, 1968), and members of the Oreville Fm as well as unamed units in the central Black Hi
Xb Metabasalt. Amphibolite, greenstone, and actinolite schist. Included are Hay Creek Greenstone (Bayley, 1972c) and part of the Vanderlehr Fm (Redden, 1968)
Xif Iron formation. Iron formation, ferruginous chert, and minor mica schist. Included are Homestake Fm, Rochford Fm and Montana Mine Fm, and iron-formation in the Keystone area and in minor areas throughout the Precambrian core of the Black Hills
Xsc Siliceous biotite phyllite, calcareous biotite phyllite, and schist. Includes the Poorman Fm (Hosted Wright, 1923) and Reausaw slate (Bayley, 1972b) which intertongue with meta basalt Xb
Xd Dolomitic marble and schist. Marble, phyllite, and calcareous phyllite. Includes the Roberts Draw Fm in the Nemo area (Redden, 1981) and part of the Vanderlehr Fm at Bear Mountain
Xqc Quartzite and metaconglomerate. In the Nemo area, map unit includes taconite conglomerate and is equivalent to the Estes Fm (Redden, 1981). In the eastern and southeastern Black Hills, map unit includes Buck Mtn quartzite (Redden, 1981) and Gingrass Draw
157 4. Unit Xsc is described as equivalent to the Poorman formation (Noble and
Harder, 1948) but it overlies unit Xqc which is equivalent to the Estes
formation (Redden, 1981) and the Montana Mine formation (Bayley,
1972b), both of which are supposed to be stratigraphically younger.
5. Parts of the Oreville formation as defined by Ratte and Wayland (1969)
appear in four different mapped units that occur at different stratigraphic
intervals: Xps, Xms, Xbs, and Xgw
If one assumes that the correlated stratigraphy is correct and then rearranges the
explanation of map units given by Redden et al. (1990) to amend the
discrepancies mentioned above, a whole new set of problems arise. If the
process is repeated, the effort becomes a shell game for which there seems to
be no unique solution. The use of the assumed correlations to interpret the
structural relationships within the Black Hills further complicates the conundrum
until one must rely on abstruse stratigraphic, structural, and tectonic relationships
to explain the overall pattern.
The root of the problem is not the total lack of stratigraphy, but the
presence of local stratigraphies that do not seem to correlate across major fault
boundaries. Furthermore, attempts to do so have become mired in terminology
such as “Bugtown” or “Oreville” formational names, and hence a genetic
relationship is implied where one does not likely exist. The abundance of similar
rock types and the disposition of previous workers to assign formational names with little corroborating evidence except for similar lithologic characteristics is
demonstrated above. Thus, a stratigraphy has evolved and has been applied
158 across the Black Hills that is inconsistent within itself, yet has remained unchallenged in the literature. Moreover, there is also clear evidence for the major structural features and relationships of the Precambrian Black Hills, yet the stratigraphic nomenclature used within the Black Hills is inconsistent with respect to the structural relationships. Hence, it is apparent that the overall stratigraphic relationships within the Black Hills cannot be integrated as easily as has been proposed in the literature.
It is clear from the map descriptions of Redden et al. (1990; Table 2-1) that either the stratigraphic correlations are not viable or the mapped structure is incorrect. There is strong evidence for the overall structural patterns shown in
Figure 2-5. It is also evident from careful reading of the literature and of Table 2-
1 that ingrained stratigraphic nomenclature used across the Black Hills must be incorrect. The similarity of rock types and depositional successions coupled with the lack of distinct marker units over much of the Black Hills has lead previous workers to construct an over-arching stratigraphic framework that does not exist in the Precambrian Black Hills.
159 APPENDIX C- Thin Section Locations
Outline map of the Mount Rushmore Quadrangle. Each dot represents a thin-sectioned sample from this study. Compare with Plate II (station numbers) for exact sample numbers. Samples from Drs. Bauer and Nabelek not shown.
160
Vita
Joseph Christopher Hill was born to Earl W. and Lorene R. Hill in Oak Ridge,
Tennessee on April 11, 1969. He is the youngest of eight children. Joe grew up on his family’s farm near the Clinch River just outside Clinton, Tennessee. He graduated from
Anderson County High School in June, 1987. In 1995, Joe decided to also pursue a degree in Geology and found his calling. He graduated with a Bachelor of Science in geology in 1996 and earned a Masters of Science degree in Structural Geology in 1999, both from the Department of Earth and Planetary Sciences, University of Tennessee-
Knoxville. Joe was married September 5, 1998 to Kristin Leigh Taylor, a beautiful brunette from Conyers, Georgia who’s love of animals and the outdoors eclipsed his own. In the Fall of 1999, Joe and Kristin moved to Columbia, Missouri to complete Joe’s graduate work in geology. While in Missouri, they had a beautiful baby boy, Aidan
Christopher, on October 11. Joe accepted a tenure-track job in 2004, defended his dissertation in 2006, and is currently a Professor of Geology at Bloomsburg University of
Pennsylvania and the family relocated to Pennsylvania. He hates to refer to himself in third person.
161